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Chapter 1 Miocene-pliocene hydrothermal ore deposits in and around the Japanese Islands
Chapter 1 Miocene-Pliocene Hydrothermal Ore Deposits in and around the Japanese Islands
1.1. General overview of metallogeny and tectonics in the Japanese Islands Before mentioning the characteristics of Kuroko and epithermal vein-type deposits in Japan, it is worthwhile to briefly describe the metallogeny, geology, geophysics, and tectonic situations of the Japanese Islands. Japan is situated along the boundary between Eurasia and the Pacific plate (Fig. 1.1). The metallogeny, geology, geophysics and geochemistry of Japan are largely controlled by this tectonic setting. The great variety of mineral deposits of Japan reflects the complex geotectonic environments. An intimate relationship exists between igneous and hydrothermal activity, which in turn reflects the plate tectonic history of Japan. Many Japanese ore deposits have produced many different metals, and they contain almost all common and useful minerals, although many deposits are small in size. Important metallic ore deposits include Besshi (Kieslager)-type (strata-bound cupriferous pyritic deposits), strata-bound Mn-Fe-type, skarn-type, Kuroko-type and vein-type. Dominant non-metallic deposits are limestone, clay, native sulfur, zeolite, silica and gypsum deposits. The deposits are divisible into three groups, based on their ages of formation: Carboniferous-Jurassic, Cretaceous-Paleogene and Tertiary-present. Carboniferous-Jurassic deposits, closely associated with submarine volcanic rocks, are of two kinds: Besshi (Kieslager)-type, and strata-bound Mn-Fe-type deposits. Besshi-type deposits are cupriferous pyritic deposits and occur mainly in metamorphic terranes (Sanbagawa, Sangun, Abukuma and Hidaka) and in some other areas (Chichibu and Shimanto; Fig. 1.2). The geological and geochemical similarities of these deposits and modern midoceanic ridge deposits (e.g., Juan de Fuca ridge, Guaymas) suggest a similar origin. For instance, the sulfur isotopic compositions of both types of deposits are equal to or higher than mantle values (generally + 1%0 to +4%0), suggesting mantle origin, perhaps modified by seawater-basalt interactions. In the Chichibu Zone, the intimate association of abundant strata-bound Mn-Fe deposits, limestone-dolomite and silica (chert) with basic volcanic rocks suggests an ocean-ridge hydrothermal origin. Jurassic-Cretaceous Besshi-type and Mn-Fe strata-bound deposits are present in Hidaka, Hokkaido (Fig. 1.2). Geochemical data and geological evidence all point to a midoceanic ridge environment of ore formation. 334S values of Shimokawa Besshi-type
Chapter 1 North American Plate
Eurasian
. ,o
Plate
.::.,y;
Sea of J a p a n -'~
Pacific Plate
•"::'. -**oh
\ Basin
~I
• :.:.
Philippine Sea Plate ' ~--..~-"~ I
TROUGH
I
/ /
PARECE VELA BASIN
/ Figure 1.1. Outline index map of the Japanese subduction zones. Thick lines with teeth are converging plate boundaries. Arrows indicate relative plate motions. Abbreviations: su, Suruga trough; sa, Sagami trough; sf, South Fossa Magna triple junction; och, Off Central Honshu triple junction; ISTL, Itolgawa-Shizuoka Tectonic Line; KSM, Kashima VLBI station (Uyeda, 1991).
deposits range mostly from +7%o to +10%o (Miyake and Sasaki, 1980), suggesting a contribution of seawater sulfate in addition to mantle source sulfur. In C r e t a c e o u s - P a l e o g e n e time many skarn-type and vein-type deposits formed associated with intense granitic activity. Granitic rocks are divisible into magnetite-series and ilmenite-series (Ishihara, 1977). Magnetite-series granitoids are present in North Honshu (Kitakami) and in the inner zone o f Southwest Honshu (San-in) Belt and ilmenite-series granitoids in the outer zone of Southwest Honshu (San-yo) belt. Metals associated with these two types of granitic rocks are distinct: Mo, Cu, Pb, Zn, Au and A g with the magnetite-series; Sn, W and rare earth with the ilmenite-series. Isotopic (Sr, S and O) data suggest that the ilmenite-series granitic m a g m a was influenced by
~ 1
Cretaceous-Paleogeneaccretionarycomplex
~
2 Cretaceoushigh P/Tme|amorphicbelt
~
3 JurassicaccreationaPjcomplex
~
4 Paleozoicaccretionarycomplex
~ 5
Continentalblock
@Cu(Bessh!subtype) ,,
c- ( . i . . o h i ~.,k,,ype)
•
~
M.
%
-/.
•
,/i~ ~ (~ ~
Sanbagawabelt
0
"~
~
S.Kitakami
HONSHU
0
100
200km
Figure h2, Distribution of the stratiform Cu-su|fide and chert-hosted Mn deposits in Japan (Sato and Ka_~e, 1996). MTL: Median Tectonic Line; TTL: Tanakura Teetonic Line; ISTL: Itoigawa-Shizuoka Tectonic Line; BTL: Butsuzo Tectonic Line.
4
Chapter 1
contamination of carbon-bearing sediments, whereas the magnetite-series granitic magma ascended from the lower crust without a significant interaction with carbonaceous matter. In the Kitakami district (North Honshu) gold vein-type (mesothermal-type) deposits and Cu-Fe skarn-type deposits occur associated with magnetite-series granitic rocks. Ore deposits associated with volcanic rocks generally exhibit polymetallic (Cu, Pb, Zn, Sn, W, Au, Ag, Mo, Bi, Sb, As and In) mineralization. Sulfur isotopic values of sulfides from these deposits are close to 0%o, suggesting a deep-seated origin of the sulfide sulfur. Clay deposits (pyrophyllite, sericite and kaolinite) are associated with both felsic volcanic rocks and ilmenite-series granitic rocks of late Cretaceous age in the San-yo Belt. Tertiary-Recent mineralization in Northeast Japan includes both epithermal veintype and Kuroko deposits (Fig. 1.3). Kuroko deposits occur only in the Green tuff region, whereas vein-type deposits occur both in the Green tuff region and in subaerial Tertiary-Quaternary volcanic region (Non-Green tuff region). K-Ar ages of formation of Kuroko and vein-type are middle Miocene (13-t- 1 Ma) and Plio-Pleistocene (3 4-2 Ma), respectively. Precious metal vein-type deposits include enrichments in Au, Ag, Hg, Te, Se, Sb, As, S, and Bi. Base metal vein-type deposits contain Pb, Zn, Mn, Ag and Cu, whereas Kuroko deposits are enriched in Cu, Pb, Au, Ag and Ba. In Southwest Japan, two styles of vein-type mineralization (Hg and Sb) formed from middle Miocene to the present. Many Hg and Sb deposits are present along the Median Tectonic Line, associated with the Setouchi andesites and ilmenite-series granitic rocks (Fig. 1.3). These different sites of hydrothermal and ore-forming activity may have resulted from the mode of subduction of the Pacific Plate. Mariana-type subduction (characterized by a steep angle of subduction and back-arc basin formation; Uyeda and Kanamori, 1979) during middle Miocene caused WNW-ESE extension, submarine hydrothermal activity, thick accumulation of bimodal (basaltic and dacitic) volcanic activity (Green tuff) and Kuroko-type formation (Shikazono and Shimizu, 1993). Plio-Pleistocene Chilean-type subduction (shallow-dipping subduction zone, E - W compression; Uyeda and Kanamori, 1979) and oblique subduction of the Pacific Plate beneath the North American Plate led to uplift and expansion of land area, subaerial hydrothermal activity accompanied by meteoric water circulation, subaerial andesitic volcanic activity and formation of vein-type deposits. Figure 1.4 shows the quantities of metals produced f¥om the metallic deposits in Northeast and Southwest Japan. This figure demonstrates that a large quantity of base metals (Cu, Pb and Zn) and precious metals (Au and Ag) was concentrated in the deposits of Northeast Japan, while they are much less abundant in the deposits of Southwest Japan. Subduction of the westward-advancing Pacific Plate under Northeast Japan was active during the Miocene, while in Southwest Japan, subduction along the Nankai Trough began more recently (5-10 Ma) and the Miocene ocean plate was inactive. These different plate motions may cause differences in quantities of sulfide ores and kinds of metals concentrated in the deposits. Lead and sulfur isotope data suggest that during subduction, pelagic sediments and altered basalt were incorporated into the magma in the island-arc trench systems of Northeast Japan and that lead and sulfur in vein-type and
Miocene-Pliocene Hydrothermal Ore Deposits
5
8, P Zn,S,~, ( £ A~) <~ Nn H£ Sb gure~ @NStS Bed~ ma~sanesee~s~s
Figure 1.3. Distribution of mineral deposits and tectonic provinces of the Neogene in Japan (Tatsumi, 1970). I, Zone of Miocene volcanism (Green Tuff region; mainly submarine); II, Zone of Miocene volcanism in the Ryukyu Arc (mainly subaeriaI); III, Zone of Neogene volcanism along the Median Tectonic Line (mainly subaerial); IV, Zone of Late Neogene folding, mainly in the Green Tuff region; V, Zone of Neogene intrusives and extrusives in the Shimanto terrane. Kuroko deposits in Northeast Japan originated from these materials. Antimony, mercury and sulfur in the H g - S b deposits in Southwest Honshu may have been derived from the shallow level of the crust under the Shimanto Group. Large epithermal gold vein-type deposits occur at major a r c - a r c junctions (Figs. 1.5 and 1.6); specifically, Chishima (Kurile)-Northeast Honshu, Northeast H o n s h u I z u - B o n i n and Southwest H o n s h u - R y u k y u . This m a y result from hydrothermal activities and mineralizations caused by intense volcanism at the a r c - a r c junctions. Hydrothermal c l a y - s i l i c a deposits (kaolinite, halloysite, sericite, montmorillonite and silica) and zeolite deposits occur in Tertiary-Quaternary volcanic regions. These deposits are distributed in areas o f epithermal gold mineralization.
Chapter 1 Miocene-Pliocene
Quaternary
Isla2d HE
i
Valcanic ' ,I
Japan
SW dapa.~n.n 0
DRI 30000
1
20000
40000
Kuroko-type deposit
6000
U/ZA
2000
Vein-type deposits
20000 100000 1 2 3
NN
Sulfur deposits
Figure 1.4. Comparison of quantities of ore deposits formed in late Cenozoic in NE and SW Japan. Weight per kilometer length of island arc (lshihara, 1978).
Quaternary sulfur deposits are distributed along the present volcanic front. Intersections of transverse faults proposed by Carr et al. (1973) and the present volcanic front coincide with the locations of clusters of the sulfur deposits (Nishiwaki and Yasui, 1974). Recently, it was found that mineralization is taking place in and around the Japanese Islands: sulfide-sulfate chimneys were discovered at back-arc depressions of the Ryukyu Arc Okinawa (Trough) and Izu-Bonin Arc (Smith Rift). The geologic settings are similar to those of the Miocene Kuroko deposits. The Ryukyu Arc belongs to the Mariana-type because its back-arc region is under extensional stress and the Okinawa Trough probably is a nascent back-arc spreading basin (Uyeda, 1991). The Izu-Bonin Arc may also be a Mariana-type, at present. However, it is likely that the arc was a Chilean type because intense epithermal gold mineralization took place at 1-3 Ma in the Izu Peninsula. Gold-rich silica precipitates at the Osorezan volcano, which is located in the most northern part of Honshu, have features very similar to epithermal Te-bearing gold vein-type deposits of the Plio-Pleistocene.
1.2. General overview and classification of hydrothermai ore deposits of Neogene age Main hydrothermal ore deposit types of Neogene age that formed in and around the Japanese Islands are Kuroko deposits and epithermal vein-type deposits. This classification is based on the form of the deposits. Kuroko deposits are strata-bound and massive in form (Fig. 1.7) and syngenetically formed on the seafloor and/or sub-seafloor environment. Vein-type deposits are fissurefilling and epigenetically formed (Fig. 1.8). Elemental association can be used to sub-classify these deposits. Major metal elements produced from Kuroko deposits are Cu, Pb, Zn, Ba, Ca, Fe, Au, and Ag. Average ore grade and tonnage are summarized in Table 1.1. Horikoshi and Shikazono (1978) classified Kuroko deposits into three sub-types: C sub-type (composite ore type),
Miocene-Pliocene Hydrothermal Ore Deposits
Figure 1.5. Three island arc junctions in the Japanese Islands (Kubota, 1994).
Y sub-type (yellow ore type), and B sub-type (black ore type), according to Cu, Pb and Zn ratios (Fig. 1.9). However, the variation in the ratio is not wide, compared with epithermal vein-type deposits. Therefore, characteristic differences in each sub-type of Kuroko deposits are not discussed here. Major epithermal vein-type deposits in Japan are base-metal type and preciousmetal type which are classified based on the ratios of base metals and Au and Ag which have been produced during the past (Table 1.2). Base-metal vein-type deposits may be divided into Pb-Zn-type and Cu-type (Otsu and Harada, 1963). However, this sub-classification is not considered here for simplicity of discussion.
T A B L E 1.1 Size and composition of ore deposits in the Hokuroku basin (Tanimura et al., 1983) Deposit name
Discovery (year)
Size (max. length x width x thickness, m)
Average ore grade Cu (%)
Zn (%)
Ph
Fe
s
(%)
(%)
(%)
1.13 2.28 1.32 2.60 I. 15 1.50 3.02 1.00 2.32 2.82 2.70 2.39
1.4 1.3 10.0 7.8 1.3 3.0 3.0 6.5 6.5 2.4 1.5 3.6
0.4 0.2 1.8 1.0 0.2 0.5 0.5
12.8 20.4 5.2 13.4 8.9 15.8 20.0
1.2
%0
0.5 0.4 0.3 1.0
7.0 24.4 14.8 3.6
14.7 23.4 6.0 15,4 10.2 18.2 23.0 8.0 8.0 28.0 17.0 21.1
Shakanai ore group (Shakanai Mines Co., Ltd.) No. I 1962 3 0 0 x 1 5 0 x 12 No. 2 1963 ?x?x7 No. 3 1963 400x 120x 6 No. 4 1963 400 x 300 x 40 No. 5 1964 350 x 70 x 13 No. 7 1965 350 x 250 x 15 No. 8 1965 430 x 170 x 40 No. II 1967 4 0 0 x l l 0 x 10
2.3 2.3 1.1 1.7 1.9 1.3 0.7
14.6 12.3 10.0 2.9 3.4 3.2 1.0
3,2 7,6 6.2 0.7 1.0 0.9 0,2
12.2 8.8 10.4 19.1 14.8 22.6 28.7
Matsuki ore group (Mitsubishi Metal Corp.) Matsuki 1964 350 × 90 x 30 Takadate 1963 150 × 150 x 30 Takadate South 100 × 70 x 15
1.9 3,74
11.8 2.00
3.4 0.80
0.89
10.1
3.32
Hanaoka-Shakanai area Hanaoka ore group (Down Mining Co., Ltd.) Tsutsumizawa [912 150 × 80 x 1 l0 Doyashiki 1916 270 x 200 x 100 Kamiyama 1919 60 x 40 x 55 Nanatsudate 1929 100 x 40 × 50 Higashi Kannondo 1935 35 x 15 x 25 O y a m a 1 Nishi 1938 60 × 13 x 30 Kannondo 1939 40 x 35 × 45 lnarizawa 1940 40 x 20 x 15 Ochiaizawa 1941 50 x 35 x 20 Oishizawa 1941 50 x 40 x 20 Oyama 2 1941 ll0x 17x30 Malsnmine 1963 600 x 400 x 110
Ezuri-Fukazawa
All (ppm)
Ag (ppm)
m m
m
Tonnage (1,000 metric tons)
441 8,946 932 677 105 60 258 34 207 90 100 30,000
0.5
57
t4.0 10. l 12.0 22.0 17.0 26.0 33.0
2.0 1.7 0.8 0.3
270 260 410 25
15.7 21.2
18.0 24.38
0.6
55
660 1,200
4.1
4.7
1.2
180
3,000
540 360
3,600 430 1,000
2,800
area
Ezuri ore group (Dowa Mining Co., Ltd.) Ezuri 1975
T A B L E 1.1 (continued) Deposit name
~. Discovery (year)
Size (max. length × width × thickness, m)
Fukazawa ore group (Dowa Mining Co,, Ltd.) Tsunokakezawa 1 1973 500 x 300 x 5 Kanayama 1976 210 × 90 x 8 Manjyaku 1979 190 x 190 x 13
Average ore grade Ca
Zn
Ph
Fc
S
AH
(%)
(%)
(%)
(%)
(%)
(ppm)
1.13 1.6 1.0
15.4 19.0 10,1
3.3 5.8 1.5
4.4 7.9 3.6
5.1 9,1 4.1
0.6
Ag (ppm)
93
Tonnage (1,0130 metric tons)
e~ I
~z
3,000
,5 g-
K o s a k a area
Uwamuki ore gnmp (Dowa Mining Co., Ltd.) No. 1 I962 1 5 0 × 1 0 0 x 14 No. 2 1965 200 × 150 x 40 No. 4 1966 350 x 1 0 0 x 17
0.6 0.8 0.8
11.5 7.8 8.3
4.2 1.8 2.8
4.1 7.7 5,5
4.7 8.8 6.3
0,7 -
Uchinotai ore group (Dowa Mining Co., L t d ) West 1959 400 x 300 x 70 East 1960 400 x 300 x 40
2.8 2.0
4.0 4.4
1.1 1.5
17.2 13.6
19.8 15.6
0.8
Motoyama ore grcmp (Dowa Mining Co., Ltd.) Motoyarna 1861 300 x 700 x 50
2.2
4.5
0.8
20.6
23.7
Furutobe ore group (Mitsubishi Metal Corp.) Yunosawa Daikoknzawa 1959 250 x 1 0 0 x 15 Daikokuzawa West $960 t70 x 70 x 40 Daikokuzawa Easl 1960 100 x 60 × 20 M a g a r i y a z a w a East 1962 200 x 80 x 15 M a g a r i y a z a w a West 1962 150 x 50 x 15
1,9 1.1 2,8 1.5 1.5
4.3 1.0 6.2 0.9 0.9
0.9 0.1 1.4 0.2 0.2
17,4 15.3 24,4
20.0 17.6 28.0
A inai ore group Ytmosawa Suehiro Daikoku Benten Yokodawara Hagoromo
0,7 4.7 2,2 1.9 2.0 1.9
tr 8.3 5,1 3.1 2.9 10.3
lr 1.7 1.3 1.2 0.7 3.5
21.8 17.6 16.1 15.7 18.3 20.0
25.0 20,2 18.5 18.0 21.0 23.0
120
130
160
930 2,580
5,240 4,000
15,000
Furutobe-Ainai area
1942 1955 1956 1957 1960 1967
200 50 180 200 150 150
x x x x × x
150 x 50 40 x 30 80 × 50 60 x 20 80 x 40 80 x 10
1.3
51
3,800
t.0 0.5 0.4 9,1
260 140 130 620
220 2,000 1,500 210
Chapter 1
10
r",, 12.912.8 2.93-049( !," ~,~,.;" 4.9~2.7\ ~ ~P "'t'~',,T/'~..,, ~[~] 3[W-] 5~b] 13
50.t < lO-50t I ~lOt o 0.II I
5.8
o 501oo 200 300km 1 ' ~ 3 " J I
I
I
I
I
(4.6)
3.6
0.6 1.1,3.7
1.1 4 0~--1.0 514 , ~1.5~1.8 ~-3.7 -4.3 Figure 1.6. Distributionand temporal and spatial relationshipof late Cenozoic gold deposits in the Japanese Islands. 1: Quartz vein-typegold deposits with little to no base metals. 2: Gold silver deposits with abundant base metals. 3: Distribution boundary of gold deposits formed during the Miocene. 4: Location of PlioPleistocene gold deposits at the actual island arc junctions. 5: Location of Plio-Pleistocenegold deposits in front of the actual island arc junctions. Numbers in the figure are K-Ar ages of epithermal Au-Ag veins (Kubota, 1994). Several sub-classifications of epithermal precious-metal deposits have been proposed: mineralogy, host-rock composition and elemental association (Lindgren, 1928), gold-silver ratios of metal weights (Ferguson, 1929; Nolan, 1933), mineral paragenesis (Nishiwaki et al., 1971), and production ratios of metals (Heald-Wetlaufer et al., 1983). Recently, epithermal gold deposits were divided into several types based on gangue minerals, and physicochemical environment of ore deposition (pH, H2S concentration of ore fluids). They are hot spring-type (Silberman, 1982; Berger, 1983a; Berger and Eimon,
Miocene-Pliocene Hydrothermal Ore Deposits
~'~A/kj~'~'+ " +~.~. L'ff~.~
.
~.
'.
J
.
~
~ .
v
~
ALTERATION
~XZ2~%8~\~,// ~ # \\%//..~.~,// ~/!
~. .
~
~,~//--,~.Stock. It Sill. . . . . Ore
AlteredB .
~,o~,p,,~
//-~ Ferruginous Chert (Fe-Mn) Barite Ore Black Ore ~ YellowOre ~
Epidote-rich Basalt 300m Acid/stuff'/,,~, ~,..--,/ ~ / ~ "
~ - ~. ..' .. ..~. ,',
11
C h l o r i t e - r i c h Basalt
O ~~" ~
+~'Y++'I:~&i ~ ,. .. .. .. .. .. .. . . ....... ~..:::iii ,o~,o
. . : . : . : ~ . ~ , ,
Clay/Mudstone Sericite/Chlorite Alteration
,~l
400m
Figure 1.7. Schematic distribution of Kuroko orebody and hydrothermally altered rocks (modified after T. Sato, 1974).
Shishimano Dacite L~
Age 0.66 l.lMa
300 , /<'/
Portal/', t
~ ~ &'~
N / \ Hishikari Lower t Andesites \
," ", - ,, ,- -
Old Hishikari'~"~"~.~• \ / / '~ ,. ,, -Yamada / ~ \
\t \ /\/'" "~ \ 1' N / \ ,
d W \/\ x,"/ x\,'\ ," \ ? ' 4 / - ~ v V ~ ' - - < ) , l ~ / \ , 0.95 ;" ,,, z \ ~ / x/\ / \ ,,"%.v,.~jc.,,v~7"-ff-~
100ML'~-,,.kL.." ~ ; - . ; ~ { , / , .
r/=/ =I~I
i~, '~
x'-"--x"-/Jf" z ""
~
I
1r
Vein
(Honko)
/
/
~-~
v
\ " \ z \ z \"~'4 1"79Ma
~
/
tl (Sanjin)
,,,
\,,
pre -Neogen~
0.78-1.05Ma Figure 1.8. Schematic northwest-trending section across the Main and the Sanjin deposits of Hishikari mine (Ibaraki and Suzuki, 1993).
12
Chapter 1
TABLE 1.2 Estimated total productions of Au, Ag and other metals and Ag/Au total production ratio (Ag/Au, by weight ratio) from the individual vein-type and disseminated-type deposits in Japan (Shikazono, 1986). Type I-A: gold-silver-rich deposits, Type I-B: base-metal-rich deposits, Type 2: disseminated-type deposits Mine
Deposit type
Au (M.T.)
Ag (M.T.)
Ag/Au
Hokuryu Sanru Numanoue Kohnomai Kitami Tokusei Taiho Teine Oe-Inakuraishi
l-A I-A I-A 1-B I-B 1-A I-A I-B 1-B
2.9 6.7 1.1 71.4 0.2 1.2 0.12 10.2 2.5
11.3 40 81.2 1219 22.2 14.7 10.2 158 109.5
3.9 6.0 71.4 17.0 11I 12.0 85 15.5 43.8
Todoroki Toyoha
1-A 1-B
5.7 2.3
209 914
37 404
Eniwa Koryu Chitose Oogane Shizukari Yagumo
1-A I-A I-A 1-A 1-A 1-B
0.71 0.76 22.8 1.6 7.4 0.4
5. I 22.2 105 35.0 5.1 70
7.1 29.4 4.7 19.4 7 175
Jokoku
1-B
-
152
-
Furokura
I-B
1.2
45
37.5
Osarizawa
l -B
6.2
251
40.5
Ani Takanosu Matsuaka
I-B I-B 1-B
1.0 0.6 1.53
32.0 3.9 62.6
30.8 6.6 40.9
Innai Hosokura
1-A l -B
1.0 2.9
400 527
400 184
Isobekoyama Handa Yatani
1-B 1-A 1-B
3.2 1 1.7
2.3 13 64
0.73 13 39
Other metals (M.T.)
Cu: 4720 Pb: 8850
Cu: 7291 Cu: I924 Pb: 17316 Zn:48100 Mn:307840 Pb: 226410 Zn: 558478
Cu: 70
Pb: 12000 Zn: 23200 Mn: 6444 Pb: 13810 Zn: 35906 Mn: 2762 Cu: 15500 Zn: 262500 Cu: 341000 Pb: 806000 Zn: 155000 Cu: 20770 Cu: 4770 Cu: 3150 Pb: 25650 Zn: 39600 Pb: 21534 Zn: 588596 Cu: 3850 Cu: 1270 Pb: 29210 Zn: 58420
Years of production
1928-1943 1925-1974 1923-1951 1917-1974 1934-1964 1930-1942 1912-1928 1932-i971 1890-1974
1903-1974 i914-1974 1929-1943 1903-1957 1936-1974 1932-1955 1918 1962 1931-1969
1941-1978
1904-I974
1885 1931-1969 1908-1950
1871 - 1953 1898-1977 1932 1919-1966 1870-1974
Miocene-Pliocene Hydrothermal Ore Deposits
13
TABLE 1.2 (continued) Mine
Deposit type
Au (M.T.)
Ag (M.T.)
Ag/Au
Other metals (M.T.)
Years of production
Mikawa
1-B
1.7
29.4
17.4
1942-1961
Sado Takatama Takahata Nebazawa Tochigi Ashio Ohito Toi Seigoshi Mochikoshi Yugashima Rendaiji
1-A 1-A 1-A 1-A 1-A 1-B 1-A 1-A 1-A I-A 1-A 1-A
57-77 28.8 3.1 1.0 0.115 3.1 1.03 18.4 13.5 4.9 2.2 5.6
1310 279.9 2.9 65 7.6 390 2.36 214 455 104 29.8 276
Cu: Pb: Zn: Cu:
9.7 0.95 65 66 125.8 2.3 11.6 34 21 13.5 50
Nawaji Shimonomoto Kishu Okinoura Takeno Kohmori Nakase Ohmidani Ikuno
1-A 1-B I-B 1-A 1-A 1-B I-B 1-A 1-B
1.5 1.3 2.2 4.5 4.6 0.123 2.6 0.3 2.1
25 80 179 4.4 91 28.0 12.8 79 403
16.7 62.5 80 1.0 19.5 228 4.9 267 194
Tada
1-B
0.0
0.2
8000
Sakoshi-Odomari OmorI Bajo Talo Fuke Okuchi Onoyama Yamaganc lsobe-Arakawa Kushikino Akeshi Kasuga Iwato
1-A 1-B 1-A 1-A 1-A 1-A 1-A 1-A 1-A 1-A 2 2 2
1.1 1.4 13.0 36.6 1.9 21 1.3 37.1 1.5 53.6 2.4 3.3 4.4
9.7 65.7
8.8 48.1
158.6 1.1 15.6 0.6 37.1 10 488 1.4 0.69 5.7
4.3 0.6 0.74 0.43 1.0 6.7 9.1 0.6 0.21 1.3
5281 1686 7857 5400
Cu: 671795
Cu: 1000 Mn: 15840 Pb: 3680 Cu: 89436
Cu: 20910 Sb: 1921 Cu: Pb: Zn: Sn: Cu: Pb: Zn:
76076 27664 152152 1521 12 2 7
Cu: 6331
1601-1970 I429-1974 1930-1976 1942-1974 1908-1950 1877-1966 1930-1952 1916-1965 1935-1976 1929-1962 1937-1972 1914-1959 1929 1971 1956-1962 1940-1974 1925-1942 1920-1949 1928-1968 1956-1966 1914-1974 1940-1973
1940-1973
1977-1982 I889-1918 -1945 1903-1973 1937 1947 1905-I974 1934-1963 i628-1955 -1970 1914-1974 1915-1974 1929-1966 1932-1980
Chapter 1
14
Cu
0
"Y" sub-type 0
0.- ~"~"~" .~s S @@
"C" sub-type
"B" sub-type
Pb
~@
•
8
•• •
\ Zn
Figure 1.9. Available data on the Cu, Pb and Zn ratio of total ore in a single unit deposit in the HanaokaKosaka district, marked with three sub-typesof Kurokodeposits (Horikoshi and Shikazono, 1978).
1983), quartz-alunite-type (Berger, 1983b; Berger and Eimon, 1983), high sulfidation and low sulfidation-type (Hedenquist, 1987), a low and high sulfur distribution and an alkali-type (Bonham, 1984, 1986), and Te- and Se-bearing types (Shikazono et al., 1990). Most of epithermal precious-metal vein-type deposits in Japan can be classed as adularia-sericite-type, and low sulfidation-type. Very few hot spring-type deposits (quartz-alunite-type, high sulfidation-type) are found in the Japanese Islands. A summary of various characteristic features of adularia-sericite type (low sulfidation-type) is given mainly in section 1.4. A few examples of hot spring-type deposits occur in the Japanese Islands. The characteristics of this type of deposits are described briefly in section 2.7. Shikazono et al. (1990) divided epithermal precious-metal vein-type deposits into Te-bearing and Se-bearing deposits. As will be considered later, Te-bearing deposits are regarded as intermediate-type of adularia-sericite-type and hot spring-type. The distinction between these two types of deposits is discussed in section 1.4. During the Miocene age, polymetallic vein-type (xenothermal-type, subvolcanictype) and gold-quartz vein-type (mesothermal-hypothermal-type) mineralizations occurred mainly in middle to western part of Japan. They are described in section 1.6.1. In section 1.6.2, Hg and Sb vein-type deposits are described. Each deposit type is distributed in a different metallogenic province (Fig. 1.3) (Tatsumi, 1970). Epithermal vein-type deposits occur in Miocene-Pliocene volcanic terrain.
Miocene-Pliocene Hydrothermal Ore Deposits
15
Polymetallic vein-type deposits occur in middle Miocene volcanic terrain in central and western Japan.
1.3. Kuroko deposits Hirabayashi (1907) defined "Kuroko" as an ore which is a fine compact mixture of sphalerite, galena, and barite. This definition can be applied to "black ore", but not to "yellow ore" or "siliceous ore" because these minerals are not abundant in these ores. Kinoshita (1944) defined "Kuroko deposit" as a deposit genetically related to the Tertiary volcanic rocks, consisting of a combination of Kuroko (black ore), Oko (yellow ore), Keiko (siliceous ore), and/or Sekkoko (gypsum ore) (Matsukuma and Horikoshi, 1970). The deposit is generally defined as a strata-bound polymetallic sulfide-sulfate deposit genetically related to Miocene bimodal (felsic-basaltic) volcanism (T. Sato, 1974).
1.3.1. Geological characteristics 1.3.1.1. Distribution Kuroko deposits occur in the Green tuff region which is characterized by thick altered volcanic and sedimentary piles of Miocene age. Distributions of Kuroko deposits and names of the representative mines are given in Fig. 1.10 and Table 1.1. Metals produced during the past are summarized in Table 1.1. Large Kuroko deposits occur in the Hokuroku district in Northeast Honshu (Fig. 1.11). Small numbers of Kuroko deposits are found in other districts such as Southwest Hokkaido, the northern part of Honshu (Shimokita Peninsula district), the middle of Northeast Honshu (Wagaomono and Aizu districts) and western Honshu (San-in district) (Fig. 1.10). It is clear in Fig. 1.10 that the distribution of Kuroko deposits is restricted in a narrow zone in the Green tuff region which was called a "Kuroko belt" by Inoue (t 969). This belt was formed by rapid subsidence under the extensional stress regime and is thought to have been a back-arc depression zone at middle Miocene age. The relationship between tectonic setting and formation of Kuroko deposits is discussed in section 1.5. 1.3.1.2. General geology, country rocks and tectonic setting A large number of studies on the general geology and stratigraphy in the Kuroko mine areas have been done. During t960-1970 many drillings were carried out by metal mining companies and the Metal Mining Agency of Japan. These data clarified geologic structure and stratigraphy of the mine areas. Many Kuroko deposits are distributed in the Hokuroku district, Northeast Honshu (Fig. 1.9). Therefore, general geology and stratigraphy of the Hokuroku district is briefly described below mainly following T. Sato (1974, 1977), Tanimura et al. (1983) and Ishikawa (1991) (Table 1.3). The lowest rock units are basement rocks, composed of phyllites, cherts, and minor sandstone probably of Paleozoic age. The oldest Tertiary formation, which is called Ohya
Chapter 1
16 oo
HOKKAIDO 200 km
HOKUROKU DISTRICT WAGAOMONO
04
DISTRICT
AIZU
HONSHU SANIN
~. ~' <~ S~~HIKOKU
~
~.~ /
JKYUSHU IG GI ( F
Q
~
~t4
DISTRIBUTION OF KUROK0-TYPE MASSIVE SULFIDEDEPOSITS IN JAPAN °.% Network, Stratiform sulfide, Gold pyrite, * °%
~,~
Pyrite, or typical Kuroko
Kuroko-typeGypsumor Barite Green TuffBeltof Japan
f;3~!~.~'~ Clusters of Kuroko Deposits
Figure 1.10. The distribution of the Green Tuff belt of Japan and the Kuroko-type massive sulfide deposits within it. Major mining districts are labeled and ore deposit clusters outlined (Cathles, 1983a).
(Menaichizawa, Sasahata), is composed mostly of brecciated andesite lavas and andesitic hyaloclastics. This formation is conformably overlain by the formation (Hotakizawa, Sunakobuchi) composed of thick sequence of basaltic lavas and tuff breccias with minor intercalations of mudstone and felsic tuff. The formation which is mostly composed of dacite lavas, tuff breccia and mudstone (Hanaoka, Yukisawa, Uwamuki formations) conformably overlies the Hotakizawa and Sasahata formations. The thickness of these formations is 300-400 m. Kuroko ore deposits occur at the upper part of this formation. White rhyolite lava domes characterized by intense sericite alteration have a close spatial relationship with Kuroko deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
N
|
~
x
x,z
,--'"
I /~
/
",,)
X~,.
/.-
Oinzan~, ~LakeTow..~ada
? /CNamanyamaI~ Furutobe ,~,. ~ ~-
/~ H~naoka! ].:/ / d~ IShakanm"/
"I
__
17
\-~'x '~Kosaka'.,~ I , _):.-~'~- ~'lt /~" " ~
-
',,/,....
r=k===W=~.-,:.\ \ 1 ~-
J
.~Somaki i" " ",~
; ( ( ./'=.... -4"
Osarizawa ~, ',~,,=
~"
•
/ t
'/~fd ]
.,.
"-,.,
O |
-'--"~ Fault ----'-1 Anticline [--'~---1 Syncline
~ ~L. ~
M. Mioc. Sed. Basin
Mioc.
Sed. Basin
10 km i
|
---] Kuroko Deposit ~
Network Vein Deposit
Principal
---1 Vein Deposit lntrusives Figure 1.11. Geologic structure and ore deposits in the Hokurokudistrict, Akita Prefecture,North Honshu (T. Sato, I974).
The formation composed of alternation of dacitic tuff and mudstone and basalt lava (Tsutsumizawa, Kagaya and Akamori formations) conformably overlies the Hanaoka, Yukisawa and Uwamuki formations. These formations are on average 150 m in thickness. The interbedded felsic tuff and mudstones (Shishigamori, Shigenai and Harukizawa formations) of middle Miocene conformably overlie the Tsutsumizawa, Kagoya and Akamori formations. The thickness of these formations varies widely, ranging from 40 m to 250 m. The younger formations (Ittori, and Tobe formations) of late Miocene to Pliocene overlie the Shishigamori, Shigenai and Harukizawa formations and are comprised mostly of mudstones, interbedded felsic tufts, and tuffaceous sandstones. The total thickness of these formations is ca. 500 m. The formations of Pleistocene unconformably overlie the
18
Chapter I
TABLE 2.3 Simplified stratigraphic column in the Hokuroku district and correlations with the Oga stratigraphy (Tanimura et al., 1983) HOKUROKU DISTRICT AGE
BLOW'S ! OGA PENINSULA M.Y ZONE
SHAKANA]MINES CO.,LTO. and AK{TA PRER GOV~
o
PLEISTOCENE -
- 5-
LATE MIOCENE
N.20 N.I9 N.)8
FunokowoE
N.I6 N.IS
Sqinokomi E (01akI F.)
*A-z~.
~-
MO-b rO-b MI-a TI
Ul-b T2182
N{ NJ N.[1
Shokona{ F. (Okuzu E)
N.S
~aolln S,dti.
_~;
TO-o
Itl6ri E
OnnogawoE (marine)
~
Tobe E
(marine)
Ealtlr,B TowedoVo~nlCl
T,.oo,o,~,.,J
BGztn~ MO-a'
k.14
MIDDLE MIOCENE
Central
J~IU¥1Um~
K]lauro I~ (marine)
N.17
IO
OOWA MININB CO., LTD.
Weslern
JI/T-W
N.2~ N.ZZ
PLIOCENE
M.M.A.j.
EARLY MIOCENE
{
~
"-~,TIUIlU~II¢IG"F'.'.-'~'. XogayeE~-,*.-.-.- Akofaorl F,o- %~
..... , HoaooknF..~. . . . . . .
Uwomukl
Su~okobuchl P
E ", : ;-~
~
T5/M3/B~
~20-
Yuk~snwa
T4/04
"
l:
'
"' "
. . . .
Ho~ok~zawa~
"
"
'
"
*.
'
'
"
. . . . .
"
~
"'"~
OLIGOCENE 25"
"60" /
MonzenG.
Ohya F. (MenoLchlzawa F.}
Toy/Ts
Oa~.ement
9aslme~!
P
~,,~;c:,lzewa F. : : : : : ' : : : : : : : : : ~~,:::::::::::::::::::::::::::::::::::::::::.::
L':::':::::::::'
Right column represents the average thicknesses and names of formations in three parts (western, central, and eastern) of the Hokuroku district, adopted by the Dowa Mining Company, Ltd.; second and, third columns from right are, respectively, the formation names adopted by the Metal Mining Agency of Japan and by the Shakanai Mines Company, Ltd., and the Akita Prefecture government; heavy horizontal and near-horizontal lines represent specific times, for example, the Menaichizawa and Sasahata formations are contemporaneous and correlate with parts of the Monzen Group and Daijima Formation at Oga.
a b o v e formations. T h e formations are basin-fill sediments, Towada v o l c a n i c s and terrace and a l l u v i u m deposits. T h e g e o l o g i c history in the H o k u r o k u district is divided into five phases as follows (Tanimura et al., 1983). ( l ) A n d e s i t e v o l c a n i s m and subsidence from near-shore to shallow-sea environments ( 3 0 ( ? ) - c a . 17 m.y. ago). (2) M a j o r subsidence, formation o f the H o k u r o k u basin and basaltic v o l c a n i s m ( 1 7 - 1 6 m.y. ago). (3) Felsic and basaltic v o l c a n i s m s and f o r m a t i o n o f K u r o k o ores ( 1 5 - 1 4 m.y. ago). (4) A c c u m u l a t i o n o f m u d s t o n e s and tufts in q u i e s c e n t sea ( 1 4 - 5 m.y. ago).
Miocene-Pliocene Hydrothermal Ore Deposits
19
(5) Differential uplifts, broad folding, andesitic volcanism and formation of smaller basins (5 m.y. ago-present). The most important geologic events at the time of Kuroko mineralizations are rapid subsidence just before the mineralization and bimodal volcanism (contemporaneous basic and felsic volcanism) (Konda, 1974). Several different hypotheses on the tectonic setting of the Kuroko mine area have been proposed. They include volcanic front of island arc (T. Sato, 1974; Horikoshi, 1975a), rifting of island arc (Cathles, 1983a), back-arc depression (Fujioka, 1983; Uyeda, 1983), and back-arc basin. In recent years, many hydrothermal solution venting and sulfide-sulfate precipitations have been discovered on the seafloor of back-arc basins and island arcs (e.g., Ishibashi and Urabe, 1995) (section 2.3). Therefore, it is widely accepted that the most Kuroko deposits have formed at back-arc basin, related to the rapid opening of the Japan Sea (Horikoshi, 1990). The summary of the bulk chemical compositions (major elements, minor elements, rare earth elements), 87Sr/86Sr (Farrell et al., 1978; Farrell and Holland, 1983), microscopic observation, and chemistry of spinel of unaltered basalt clarifies the tectonic setting of Kuroko deposits. Based on the geochemical data on the selected basalt samples which suffered very weak alteration, it can be pointed out that the basalt that erupted almost contemporaneously with the Kuroko mineralization was BABB (back-arc basin basalt) with geochemical features of which are intermediate between Island arc tholeiite and N-type MORB. This clearly supports the theory that Kuroko deposits formed at back-arc basin at middle Miocene age.
1.3.1.3. Age of mineralization The age of Kuroko mineralization can be estimated from (1) K - A r ages of igneous rocks associated with Kuroko deposits and (2) foraminiferal assemblages in mudstone directly overlying Kuroko deposits. Many K - A r ages of igneous rocks in Kuroko mine area in Hokuroku district yield 11-16 Ma. Ohmoto (1983) considered based on these data that the age of Kuroko formation was 11-16 Ma. However, almost all of igneous rocks associated with Kuroko deposits were hydrothermally altered after the mineralization (Shikazono et al., 1998). Therefore, it is likely that the K - A t ages of altered igneous rocks give younger ages than that of Kuroko mineralization. In contrast to K - A r ages, foraminiferal ages yield more accurate ones. Horikoshi (1987) investigated the foraminifera1 assemblage in mudstone overlying Kuroko deposits in the Shakanai area of Hokuroku district and concluded that the Kuroko mineralization age is 15-16 Ma. However, foraminiferal age data obtained by Horikoshi (1987) are scarce. Yoshida and Yamada (2001) compiled K - A r age of igneous rocks in Hokuroku district (which are considered to have formed simultaneously with Kuroko deposits) at 12.7 Ma. Considering uncertainty of foraminiferal and K - A r ages it seems reasonable that the Kuroko deposits in Hokuroku district formed in 14-12 Ma (more likely 13.612.7 Ma).
20
Chapter 1
However, several small Kuroko deposits (e.g., Yunosawa in Hokuroku district, Kuroko deposits in Hokkaido) occur in the formation younger than middle Miocene age (16-14 Ma), suggesting younger ages (12-13 Ma). 1.3.1.4. Metals enriched and metal ratios Many elements are concentrated in Kuroko deposits. Metals and minerals recovered from the ores are Cu, Pb, Zn, Fe, Au, Ag, S, BaSO4 and CaSO4. Size, average ore grade and tonnage of representative ore deposits are summarized in Table 1.1 (Tanimura et al., 1983). Total tonnage concentrated in the Hokuroku district is Cu, 1.2 x 108 ton; Zn, 2.5 x 106 ton; and Pb, 0.8 × 106 ton, respectively. Tonnage of BaSO4 and CaSO4 is given in Table 1.4. Figure 1.9 shows the proportion of Cu, Zn and Pb contents of Kuroko ore (Tatsumi and Ohshima, 1966; Horikoshi and Shikazono, 1978). Horikoshi and Shikazono (1978) divided Kuroko deposits in the Hanaoka-Kosaka area of Hokuroku district into three sub-types based on the ratio of Cu to Pb and Zn which increases in order of the B (black ore), C (composite ore), and Y (yellow ore) sub-types (Fig. 1.9). Characteristic features of these three sub-types were summarized by Horikoshi and Shikazono (1978) and are briefly decribed below. The Shakanai No. 1 deposit in the Shakanai mine is a good example of the B subtype (Kajiwara, 1970a). The ore of the B sub-type deposits consists of predominantly of galena and sphalerite with lesser amounts of chalcopyrite. Ore deposits of this sub-type are usually not directly associated with dacite lava dome. However, it is known that domeshaped dacite occurs below some of this sub-type deposit (Kajiwara, 1970a; Tanimura et al., 1974). Total ore quantity of a single unit deposit is generally small, about one million tons. C sub-type deposits are often called typical Kuroko deposits. Sato (1970) and Horikoshi (1976) published the schematic sections of Kuroko deposits referring to the general geology of this sub-type. The major Kuroko deposits belong to this sub-type. The largest Kuroko deposit is the Doyashiki deposit in the Hanaoka mine which belongs to C sub-type. The total ore quantity may be more than 10 million tons. The second largest deposit of Kuroko deposits is the Motoyama deposit of this sub-type. About 7 million tons of
TABLE 1.4 Estimated total amount of barite and sekko (gypsum + anhydrite) Shikazono, t983) BaSO4 (105 t) Kosaka
Fukazawa
Uchinotai Uwamuki Motoyama Tsunokakezawa Manjaku Kanayamazawa Matsumine
13. l 5.7 I 1.8 10.6 2.6 1.5 12
CaSO4 (105 t) Wanibuchi Fukazawa Yokota Motoyama Hamago
16 1-2 14 10
Miocene-Pliocene Hydrothermal Ore Deposits
21
ore were mined. The ore contains 2.3 million tons of black ore, 1.1 million tons of yellow ore and 3.7 million tons of siliceous ore. This means that ore deposit consists of roughly equal quantities of black, yellow and siliceous ores. A part of geology of the C sub-type is exhibited strikingly in the abandoned open-pit of the Motoyama deposit. The Uchinotainishi deposit of the C sub-type was described by Horikoshi (1969). The hydrothermal activity responsible for the mineralization of the C sub-type deposits was mostly preceded by the uplift of lava dome and the subsequent steam explosion (Horikoshi, 1969). Some of the Kuroko deposits consist predominantly of pyrite containing a small amount of chalcopyrite. The ore deposits consisting predominantly of pyrite, either with an economical value of chalcopyrite or not, are called the Y sub-type deposits, which occur above dacite lava dome or lava flow, while copper-poor deposits occur mostly in pyroclastic rocks and are associated with a large amount of gypsum. The Matsumine deposit in the Hanaoka mine is typical of the Y sub-type. The Matsuki and Takadate deposits in the Matsuki mine are also classed as this sub-type (Kuroda, 1978). Many pyrite-rich ore bodies
1oJG>. t omk2
°m/ m
~
~ U c h i n o t a i - n i U
-200m
-204~ •/
\
m 0
-100 m
-200 m
Figure 1.12. Distribution of two different sub-types of Kuroko deposits in the Kosaka district, Akita Prefecture. "Y" sub-type deposits have not yet been discovered in the area. The top pre-Tertiary basement is contoured showing some depressed structures (Horikoshi and Shikazono, 1978).
Chapter 1
22
associated with a large amount o f g y p s u m ore were mined in the Hanaoka mine. It seems likely that these ore bodies are composed o f several unit deposits o f this sub-type. Figure 1.12 shows the areal distribution o f the B and C sub-type deposits in the Kosaka district. The Y sub-type deposits have not yet been found in the district. It appears that two zones characterized by the distribution of each sub-type deposit are distributed north-southernly in the Kosaka district as well as in the Hanaoka district (Fig. 1.13). Pyroclastic rocks in the Kosaka formation, in which all deposits occur, become thicker to the east, and probably moved from the eruptive centres to the east (Horikoshi, 1969). These types of evidence may indicate that the sea at that time became deeper to the east. Figure 1.12 shows also the top of the pre-Tertiary basements. Ore deposits, either B or C sub-type, occur above the crater-like depressions o f basements. The Shinsawa deposit is the sole example o f B sub-type in the midst of the H a n a o k a - K o s a k a district, so-called Hokuroku basin (Fig. 1.13). The Tsunokakezawa deposit in the Fukazawa mine and ore deposit in the Ezuri mine are also the B sub-type.
I
.( \
/,Ioo..
g ; ',,.l_..~"
Tsutsumlzawo
\
J
N
Ok
X .\_,oo.
:\
',,x
/-"
m
Bsub,ype
"C" sub-type
\ -.,oo,,
/
"/~ ~
~
J
G ShokooaiNo,tf
0
1000 m
Figure 1.13. Distribution of three different sub-types of the Kuroko deposits in the Hanaoka district. The top of MI mudstone is also shown to visualize the structure of country rocks (Horikoshi and Shikazono, 1978).
Miocene-Pliocene HydrothermaI Ore Deposits
23
The Y, C and B sub-types roughly correspond to types 1, 2 and 3 as defined by Urabe (1974a), who classified Kuroko deposits based on hydrothermal alteration and ore mineral assemblages: type 1, kaotinite-pyrophyllite~diaspore-type; type 2, sericitechlorite-type; type 3, sericite-chlorite-carbonate-type. Hydrothermal alterations in the Kuroko mine area are described in section 1.3.2. Most large Kuroko deposits belong to type 2 (or C-subtype). Type 1 occurs mostly in Northeast Honshu and Hokkaido. Type 3 deposits are distributed in middle Honshu and Southwest Honshu (San-in district). Most of the previous studies have been carried out on the deposits in the Hokuroku district. A summary of the mineralogical and geochemical characteristic features of Kuroko deposits in this district is given below (sections 1.3.2 and 1.3.3).
1.3.2. Mineralogical characteristics
1.3.2.1. Metal zoning, and ore and gangue minerals
Typical Kuroko deposits (C sub-type according to Horikoshi and Shikazono (1978)) are usually composed of gypsum ore, siliceous ore, yellow ore, black ore, barite ore and ferruginous chert ore in stratigraphically ascending order (Fig. 1.7). The main constituent minerals in each ore are as follows: gypsum, anhydrite, Mg-chlorite (gypsum ore), quartz, pyrite, chalcopyrite (yellow ore), sphalerite, galena, pyrite, barite, chalcopyrite, tetrahedrite-tennantite, bornite, electrum (black ore), barite, quartz (barite ore), microcrystalline quartz, hematite (ferruginous chert ore). Figure 1.14 shows the distribution of minerals in each ore zone (Matsukuma and Horikoshi, 1970). The occurrence of ore minerals in Kuroko deposits was described in Shimazaki (1974), Matsukuma et al. (1974) and Urabe (1974a). Sphalerite has been studied by many workers and has been used to restrict the chemical environment of ore deposition. Iron contents of sphalerite from replacement siliceous ores and fissure-filling vein ores are much higher than those from layered ores (black ore and yellow ore) (Takahashi, 1963; Sato, 1969; Urabe, 1974b; Urabe and Sato, 1978). That is to say, iron contents of sphalerite generally decrease stratigraphically upwards in a single unit of ore deposit (Urabe and Sato, 1978). Urabe (1974b) and Urabe and Sato (1978) explained this trend by increasing of oxygen fugacity due to the mixing of hydrothermal solution with ambient cold oxygenated seawater. However, it is also likely that this trend was caused by the decreasing of temperature towards stratigraphically upwards. Generally, iron contents of sphalerite buffered by iron silicates such as chlorite decrease with decreasing of temperature in active geothermal systems (Fig. 1.15) (Hayba et al., 1985). Chlorite is common gangue and alteration minerals in Kuroko deposits. Therefore, this process seems plausible. Iron contents of sphalerite are different in layered ores in different ore deposit and different sub-types of Kuroko deposits. Iron contents of sphalerite from the B sub-type deposits (Uwamuki No. 4, Shakanai No. 1, Ezuri, and Fukazawa deposits) show wide range but generally less than 0.2 wt% (Ono and Sato, 1995). The average value, however, is probably lower than the C sub-type deposits (e.g., Uchinotai deposits). This may
Chapter 1
24
Mineral
Yellow ore Siliceousl ore Siliceous PowdeP/ Bedded PyCpSp ,to J YO YO E}O
Black ore .... GaFz Mononlne BO~nite BO raI~BO BO
Pyrite
Barite ore
Ferrugineus quartz
I
Chatcopyrite Bornits
u
Idaits
Fukuchilite Chalcocite
LII L
Covelline
,, ,
,L.
ass
Betlchtlnite Fehlerz
m
m m
Bournonite ,.
!,
., ,
Enargite Luzonite
immm mm
Argentite
l W
INm(m mmm me m m m m
Stromeyerits Jalpaite Polybasite Pyrareyrite Galena Sohalerite
IIIIIII i
m ~
III
Hi-disulphide
I
mm
!P
Mfflerite i iii
Bi,mlnerals iiiii
Electrum Silver
?
Sulphur Magnetite
4IlL
Hematite
i
Quartz I
III
I
Carbonates i ill
ii
III
Clays Barite
m
Figure 1.14. Schematic diagram showing mineralogical changes in various kinds of ores of Kuroko deposits (Matsukuma and Horikoshi, 1970).
suggest that the C sub-type deposits formed at higher temperatures and close to the volcanic centre and B sub-type deposits are distal type and formed at lower temperatures. Tetrahedrite-tennantite composition varies widely in Kuroko deposits (Yamaoka, 1969; Yamaoka and Nedachi, 1978a; Yui, 1971; Horii, 1971; Shimazaki, 1974; Kouda, 1977; Shikazono and Kouda, 1979; Ono and Sato, 1995; Ishizuka and Imai, 1998).
Miocene-Pliocene Hydrothermal Ore Deposits
25
0.5 ÷
0.4 0.3 0.2 03
0.1
14.
0 -0.1 0
,,-I
-0.2 -0.3 -0.4 -O.5 200
220
240
260
280
300
HomoganizationTemperature°C Fig. 1.15. Diagram showing the homogenization temperature of fluid inclusions vs. the iron content of the host sphalerite growth zone for sample locality NJP-X on the OH vein. The line shows the predicted iron content of the sphalerite if the sulfur fugacity of the system had been buffered by the triple point - - Fe-chlorite (daphnite), pyrite, hematite (Hayba et al., 1985). Generally, tetrahedrite-tennantite composition from Kuroko deposits is characterized by high Zn content, low Fe content, high Cu content, and low Ag content compared with those from vein-type deposits in Japan (Fig. 1.16). Rarely, it contains Hg up to 1 wt% (Ishizuka and Imai, 1998). Tetrahedrite-tennantite composition varies widely in a single orebody. For instance, Kouda (1977) analyzed tetrahedrite-tennantite from Fukazawa-Tsunokakezawa deposit and showed that the Fe and Zn contents are in a range of 0-5.5 wt% and 4 . 5 - 1 0 wt%, respectively. Wide compositional zoning and heterogeneity in a tetrahedrite-tennantite grain are c o m m o n (Yamaoka, 1969; Yui, 1971; Yamaoka and Nedachi, 1978a). Positive correlation between Zn contents of coexisting tetrahedrite-tennantite and sphalerite exists (Shikazono and Kouda, 1979). This relation can be explained in terms of the following reaction.
[(Cu, Ag)10Zn2(As, Sb)4Sl3]tet -]- (FeS)sp = [(Cu, Ag)10Fe2(As, Sb)4Sl3]tet +
(ZnS)sp (1-1)
where sp = sphalerite, and tet = tetrahedrite-tennantite. Electrum occurs mainly in the black ore zone dominantly composed of sphalerite, bornite, galena and barite. It is common in B sub-type such as the Ezuri and Fukazawa deposits. It occurs in brown compact black ore consisting of barite, galena,
Chapter 1
26 AG WT%
FE WT%
60
8
7
50
6 40
t'~: A
~
';',
A
5 4
30
3
20
2 10
1 f
10
20
30
40
1
50 CUWT%
2
3
4
5
6
7
8
9 10 11 12 ZN WT%
SB W'P/o 35
C
30 f..J t'. 25
D ",
A
,~
20
~ C ~'D
15
is
2o
:Is
AS WT%
Figure 1.I6. Chemical composition of tetrahedrite-tennantite (Shikazono and Kouda, 1979). A: Au-Ag vein-type deposits, B: Kuroko deposits, C: Taishu Shigekuma Pb-Zn vein-type deposits, D: Skarn deposits (Kamioka). sphalerite, tetrahedrite-tennantite, chalcopyrite, bornite, and Ge-bearing minerals (argyrodite), stromeyerite, pearceite, and mckinstryite (Ono and Sato, 1995; Ishizuka and Imai, 1998). Electrum is often associated with pearceite and/or tetrahedrite-tennantite (Ishizuka and Imai, 1998). Exceptionally, electrum occurs abundantly in the siliceous ore of the Nurukawa deposit (Yamada et al., 1987). The mode of occurrence of electrum was described by Matsukuma (1985). Chemical compositions of electrum from Kuroko deposits were summarized by Shikazono (1981) and Shikazono and Shimizu (1988a). The Ag content of electrum from Kuroko deposits varies widely from 4.7 to 89.4 atomic% (Fig. 1.17). Electrum with low Ag
Miocene-Pliocene Hydrothermal Ore Deposits
27
40
u
0"
30
LL
20
10
0
20
40
60
80
100
Nag Figure 1.17. Frequency histogram for the Ag content of electrum from Kuroko deposits in Japan (Shikazono and Shimizu, 1988b).
content occurs in siliceous and yellow ores. For instance, electrum in siliceous ore from the Nurukawa deposit contains about 20 Ag atomic%. Electrum with high Ag content occurs in black ore. Electrum in the brown ore occurring in upper part of black ore contains Hg up to 11 atomic% and Ag contents are positively correlated to Hg contents (Ishizuka and Imai, 1998). Native silver is found in bornite-rich black ore (Matsukuma and Yui, 1979; Matsukuma, 1985) and it is thought to be secondary mineral. Compositional zoning in electrum grain is common (Shimazaki, 1974; Imai et al., 1981). The Ag content of rim of electrum grain is higher than that of core. Although Ag content varies widely, it is generally lower than that of epithermal vein-type deposits. Although analytical data on bornite are few, some data show high Ag contents (max. 1.45 wt%) (Matsukuma, 1985). Pyrite is the most abundant ore mineral. It occurs as euhedral, framboidal, and colloform forms. Abundance of framboidal pyrite increases stratigraphically upwards. Colloform pyrite contains appreciable amounts of As and Cu (Nakata and Shikazono, unpublished), whereas these contents of euhedral and framboidal pyrite are less than the detection limit of an electron microprobe analyzer. Ishizuka and Imai (1998) found that the As content increases toward outer rim and reaches up to 5 wt% in the rim of colloform pyrite from the Fukazawa deposit.
Chapter 1
28 1.0.
~2:~?~~2 o' 002 $2
o
oI 03 of 04
0-8-
05
0"6" o6
oO6
=,
06 I.L
06
0606
O.Z,-
06
0-2
9oo-9 08 " ~ o 8 8 o8
o9 0"8
1~0
1'.2
1J4
o707 1:6
I ~8
2:0
A~ in 4 ( A I , S i )
Figure 1.18. Variation of Fe2+/(Fe2+ + Mg) and tetrahedral AI of chlorite from hydrothermal ore deposits: Japanese Neogene Cu-Pb-Zn vein-type (open circle) and Kuroko deposits (solid circle). Localities: 1 Ashio, 2 Yatani, 3 Toyoha, 4 Kishu, 5 Sayama, 6 Mikawa, 7 Furutobe, 8 Hanaoka, 9 Wanibuchi, 10 western Bergslagen (Shikazono and Kawahata, 1987). Dominant gangue minerals in Kuroko deposits are quartz, barite, anhydrite, gypsum, chlorite, sericite, and sericite/smectite. Morphology o f quartz changes from euhedral in the centre to the irregular in the margin of the deposits (Urabe, 1978). No amorphous silica and cristobalite have been found. Quartz is abundant in siliceous ore, barite ore and tetsusekiei ore. Minor amounts o f M g - m i n e r a l s (talc, M g - c h l o r i t e ) occur in sekko ore. Chlorite occurs in sekko ore and it contains high amounts of Mg (Fig. I. 18). Kuroko deposits are characterized by large amounts of sulfate minerals (barite, anhydrite, and gypsum). Estimated total amount of barite and sekko (gypsum ÷ anhydrite) from individual deposit is shown in Table 1.4. Sr contents o f gypsum, anhydrite and barite
Miocene-Pliocene Hydrothermal Ore Deposits
29
Matsumine
6O
c- 10-
<
0 ..Q
5
E Z .
.
.
.
.
500
1000
.
.
.
1500 Sr
2000
con tent
(ppm)
Shakanai
t~o
t'- 10
<
0 r.~
5
E Z
2o o Sr c o n t e n t (ppm)
Fukazawa
cO (l) cO >,
c- 10
<
0
$ ..Q
5
Z ....
N n .F1. ,
500
•
-
,
I000
•
.r]
-
,
.
.
.
.
t
~500 2000 Sr c o n t e n t (ppm)
Figure 1.19. Strontium contents of anhydrites from the Matsumine, Shakanai, and Fukazawa deposits (Shikazono et al., 1983).
are 150-2,000 ppm (gypsum, anhydrite), and 0.3-3 wt% (barite), respectively (Shikazono et al., 1983) (Fig. 1.19). Mixed layer clay mineral (sericite/smectite) is found in Kuroko ore bodies and altered dacitic rocks underlying the ore. This mineral is thought to have formed by the
30
Chapter 1
interaction of sericite with hydrothermal solution whose pH decreased by the deposition of sulfides by the reaction of MC12(aq) + H2S -+ MS + 2H + + 2C1- (where M is metal such as Zn and Pb) (Tamura, 1982). Tamura (1982) thought of the following mechanism of formation of mixed layer clay mineral (sericite/smectite). Mg 2+ and Ba 2+ in hydrothermal solution may play an important role to form the interstratification of the minerals. From the experiment in which Li + has caused the formation of the interstratified structure, Mg 2+ with most similar ionic radius (0.65 A) to Li + (0.60 ,~) may accelerate the interstratification. Tomita and Sudo (1981) and Shimoda et al. (1974) have synthesized the mixed layer minerals from sericite reacted with LiNO3 solution. Li + gets into the unoccupied site in the octahedral layer of dioctahedral muscovite (sericite). To retain electroneutrality in muscovite structure, interlayer K + is easily released. Potassium is held strongly adjacent to interlayer regions where the potassium is replaced, i.e., the two regions will alternate. This is interstratification. Reichenbach and Rich (1968) have revealed that interlayer K could be removed almost completely from muscovite by using 0.1 N BaC12 solution at 120°C. This K - B a exchange may induce to form the interstratification. To verify the above interpretation, the analysis of Ba and Li in sericite/smectite is required. Preliminary analysis of sericite/smectite shows high Ba content (Ogawa and Shikazono, unpublished).
1.3.2.2. Hydrothermal alteration Several studies identified the hydrothermal alteration halo in the dacitic rocks surrounding Kuroko deposits (Fig. 1.20) (Shirozu, 1974; Utada et al., 1974, 1981; Utada, 1980; Ishikawa et al., 1976; Izawa et al., 1978; Date et al., 1983; Urabe et al., 1983; Ishikawa, 1988; Marumo, 1989; Inoue and Utada, 1991; Shikazono et al., 1995). For example, Date et al. (1983) recognized the following alteration zones in the Fukazawa Kuroko mine area of Hokuroku district from the centre (near the orebody) to the margin: (1) sericite-chlorite zone (zone III in Figs. 1.20-1.22) characterized by quartz + sericite -t- Mg-rich chlorite; (2) montmorillonite zone (zone II in Fig. 1.20) characterized by Mg.Ca-type montmorillonite + quartz -t- kaolinite 4- calcite 4- sericite -tFe-rich chlorite; and (3) zeolite zone (zone I in Fig. 1.20) characterized by clinoptilolite + mordenite + Mg.Na-type montmorillonite 4- cristobalite 4- calcite or analcime + Mg.Na-type montmorillonite + quartz + calcite -t- sericite -t- Fe-rich chlorite (Fig. 1.20). Kaolin minerals (kaolinite, dickite, nacrite), pyrophyllite and mica-rich mica/smectite mixed layer mineral occur as envelopes around barite-sulfide ore bodies in the footwall alteration zones of the Minamishiraoi and Inarizawa deposits, northern part of Japan (south Hokkaido) (Marumo, 1989). Marumo (1989) considered from the phase relation in A1203-SiO2-H20 system that the hydrothermal alteration minerals in these deposits formed at relatively lower temperature and farther from the heat source than larger sulfide-sulfate deposits in the Hokuroku district. Date et al. (1983) found the existence of an Na20-depleted dacite mass with a lateral dimension of 1.5 x 3.0 km immediately below the ore horizon (Figs. 1.23 and 1.24) and the mass is useful indicator of exploration of Kuroko ore deposits. This Na20 depletion is considered to be due to the destruction of plagioclase attacked by potassium-
31
Miocene-Pliocene Hydrothermal Ore Deposits
TKIS;
/
I
N
TKI?9
f
D-3
I
o
I
TKt$3
®
/
/ I
/
/
\
/
e
TKII9
! I iK,ao
TKI78
•
t
\
zoNE,
0
e
\
J
/
DE 4
g
/I
;rK139 ~
I
,5~\ g~,oa
TK208 e
/
/ ZONE
II'l
tK3s III
/
1.0kin
TKI81 ®
/
/I 1
I
ZONE I
I I I
HO7
TKIgO 0
0.5
i
KA1
OES
I
TK177
TKt98
e
TK 20g e
\
®
\
~'~
\ \\\
o~~
TK200
e
'
I \ T~a,
\
LEGEND ORE
DEPOSITS
(3~i
."'''4
...........' ORE ZONE
~
, , ~ '~" - ' ~ " ~ .
T,.~.. TKt/~
~K,o3)
~"'
ALTERED DACITE I N FOOTWALL
....
•
ZONE I
4)
ZONE
II
O
ZONE It"
•
ZONE IU
•
ZONE
IV
TK207 e
TKI&~ e
TK 1~,6 e
ZONE BOUNDARY
Figure 1.20. Zoning map of alteration minerals in unit D3 around the Fukazawa deposits (Date et aI., 1983).
rich hydrothermal solution and the formation of K-sericite in discharge zone. It is thought that potassium is added from ascending hydrothermal solution and sodium in dacite removed to hydrothermal slution. Hashiguchi et al. (1983) also showed that the variation of Na20 content in the footwall rock is particularly useful for detailed exploration based on the large number of chemical analyses of footwall rocks and statistical treatment of the geochemical data. Singer and Kouda (1988, 1992) confirmed based on the statistical analysis of the distribution of minerals and bulk compositions of altered rocks in the Hokuroku district
Chapter 1
32 N
om
S
=
o
°T 1
-:7-
° T2
LEGEND
T1, MI.T2, M2 :ROCK UNITS IN HANGINGWALL
0.5
i
:ZONE I
~:ZONE
U ~nd II
~:ZONE
Ill
:INTRUSIVE OOLERITE
~:ZONE
IV
: BASALT LAVA
................ :ZONES BOUNOARY
:KUROKO DEPOSITS
~FF~:Na POORPART IN FOOTWALL
T3, D3, T4,04 :ROCK UNITS IN FOOTWALL :INTRUSIVE DACITE
{ ~
0
•
1,0 km
Figure 1.21. Zoning of alteration minerals in the Fukazawa area. The location of the profile line is shown in Fig. 1.20 (Date et al., 1983).
HANGINGWALL I u~ ~
ZEOLITE
11
It'
Ill
IV
ANALCI ME MOROENITE
MONTMORILLON I T E
Ca-MONT.
No-MONT.
0
==~ S~RICITE
~u
-~ i
0
CHLORITE
PLAG IOCLASE HYDROTHEIIMAL ALTERATION DI A G E N E ~ L,~
FOOTWALL I[ u') c= ~ t~ Q:
ZEOLITE
MORDEN I T E
NtONTMORILLON1TE
Co-MONT.
~
SERICITE
~
CHLORITE
III
Mg-rlch
PLAGIOCLASE HYDI !OTHERMAL
ALTERATION DIAGENES I S )EPLETIVE
ADDITIONAL
ELEMENTS
ELEMENTS
CoO
Na20 SiO2
CoO
No20 CoO, FeO
IMgO !FezO3
Figure 1.22. Summary of alteration minerals and zoning around the Fukazawa deposits (Date et al., 1983). Legend: unbroken line = present or formed in considerable quantity; dashed line = present in small amounts, or uncertain formation; definition of zones by Date et al. (1983).
Miocene-Pliocene Hydrothermal Ore Deposits
i
~FIAe..
"
/J~
. %.
I
;<-'~
~,
•
.
.'#
.~. ..~° •
•
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. ~
o
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~
.
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o
.
.
o
o,/7/
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%
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.
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33
o
1".
•
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o
°
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~
o
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°
- ~ "~;~//~.~7~,///~ ' ../
..,
/5"
• / .
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o
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".
°
~°
~-
"~.'~No%
o 0~.
o
•
~
•
°
~,o
i
o
o o
~'\ - ~ - ~ d . 0. . . .
~,,,.~o
~
• --
•
•
"~-2- FOoIw.II Expo .... (4 l
V
KUROKOOeposit
X /
,~lm01e,
/ /
t--7 I
Oril, hole
[_...v "7, token
o 13 or fewer sornples token
.~.'~ ~ % .
~-~r-.<~z.
~
• \
~, f
t
_"'~1 ~ > ' \ -
e~
~
,~:~.;'-. ..
o o
"
I No 5
~-,,g.,~./~%Z; e
x
/
/
°
• •
"'v06.--I~°.. I__/' o \ ~''Z-'~III/Xi .~./ '~'~
o
"
•'~Z'Z.'~';¢L/i° • \
o
" °
°
". 0100
soo
,ooo~
F i g u r e 1.23. Distribution m a p o f low N a 2 0 a n o m a l i e s in the K o s a k a a r e a ( D a t e et al., 1983).
that Na20 depletion, as well as the presence of sericite, anhydrite and gypsum near the orebody, is a useful indicator for the exploration of Kuroko deposits. Figures 1.25 and 1.26 show the map of probabilities of centres of Kuroko deposits based on the data by Singer and Kouda (1988). Green et al. (1983), Urabe et al. (1983) and Matsuhisa and Utada (1993) showed that the analysis of whole-rock 6180 values is a useful method of exploration for volcanogenic massive sulfide deposits. They found that concentric zoning patterns of the whole-rock 8180 values and of alteration minerals are particularly well developed in the footwall volcanic rocks in Kuroko mine area. For instance, 8180 values of igneous and sedimentary rocks are +16.9-4-2.7%o in the zeolite zone, +11.1+2.5%o in the montmorillonite zone, and +6.7 4- 1.3%o in the sericite-chlorite zone in the Fukazawa area (Green et al., 1983) (Figs. 1.27 and 1.28). Figure 1.29 shows the 6180 variations as functions of temperature and water/rock ratio (Green et al., 1983). This calculated result indicates that the spatial variation of 8180 values of the footwall rocks and alteration
Chapter I
34
. .
.
TWNOKAKEZAWA No2
.
$? /
LEGEND Low Na, 0 Zonr (Naz0<0. 3 2 1 )
SARUMA Vrln
KUROKO 0.porit Drill hole 15 or more w r n p l * ~ taken 0
./
0
.
/
4 /
!4 OT f.**,
sornpt..
y/q./. >.
token
M 0
100
500
1000 J
Figure 1.24. Distribution map of low Na20 anomalies in the Fukazawa area. The distribution of green dacite is also shown (Date et a]., 1983).
mineralogy can be interpreted as a result of interaction between the rocks and seawater at different temperatures (25-200°C for the montmorillonite zone and 200400°C for the sericite-chlorite zone) under water-dominated (e.g., waterlrock ratio > 1) conditions and most of the hydrogen isotope data of the whole-rock samples from the Hokuroku district (6D = -34%0 to -80%0) can be explained by this interpretation (Green et al., 1983). However, the variation of of altered rock depends not only on waterlrock ratio and temperature, but also on the degree of mixing of hydrothermal solution and cold seawater. This mechanism is considered in section 1.4.2. Ohmoto et al. (1983) found that 6 1 8 0 values alteration zones and major and base metal contents of footwall dacite along the section in the Fukazawa area correlate with each other (Fig. 1.30). Shikazono et al. (1998) found that carbonates are common alteration minerals in the Uwamuki mine area of Hokuroku district and carbonate alteration superimposed on chlorite alteration. They showed that the mode of occurrences and the Mg/(Mg
+
Miocene-Pliocene Hydrothermal Ore Deposits
I
N 5 km
Probability ~ ~7
P
[~
0.001<=P<=0.1
~
35
FURUTOBE
/
0.1
P>0.5
1
Kuroko Deposit Area considered ~
~
i
UCHINOTAI SHAKANAI
ODATE
TOWADA MINAMI
Figure 1.25. Map of probabilities of centres of Kuroko deposits based on sodium depletion, sericite, and gypsum plus anhydrite(Singerand Kouda, 1988).
Fe) ratios of magnesite and dolomite occurring in hanging wallrocks are useful in the exploration for concealed volcanogenic massive sulfide-sulfate deposits. 313C and 3180 of carbonates from the Kuroko mine area are plotted in Fig. 1.31. 313C and 3180 data lie between igneous (3~3C = -7%o, 3180 = +8%0) and marine carbonate value (313C = 0%0 and 3180 = +20%o). This indicates that magnesite and dolomite formed due to the interaction of hydrothermal solution with the biogenic marine carbonates. Dolomite and magnesite data are plotted close to marine carbonate values, suggesting that they formed in the central zone close to the ore bodies due to the interaction of hydrothermal solution with the biogenic marine carbonates (Fig. 1.31). 313C and 3180 of manganoan calcite in altered basalt directly overlying Kuroko orebody are close to igneous carbonate values, suggesting they formed from ascending hydrothermal solution at discharge zone.
36
Chapter 1
Figure 1.26. Map of probabilities of centres of Kuroko deposits based on sodium depletion in the Hokuroku district (Singer and Kouda, 1988).
Superimposed alterations are common in the Kuroko mine area (Inoue and Utada, 1991). For example, K-feldspar, kaolinite, alunite, pyrophyllite and diaspore alterations cut chlorite alteration, indicating that they formed later than chlorite alteration (Inoue and Utada, 1991). Inoue and Utada (1991 ) thought, based on detailed descriptions of the hydrothermal alterations in the Kamikita mine area, North Honshu, that hydrothermal alterations in this district started from 13 Ma and ended at 3 - 4 Ma. Pyrophyllite and diaspore alterations were reported from several Kuroko deposits, although they are not common (Urabe, 1974a). This type of hydrothermal alteration is thought to have occurred at a later stage than the hydrothermal alterations associated with Kuroko mineralization (sericite, chlorite, and zeolites) (Utada, personal communication, 1995). As well as felsic volcanic rocks, basalt occurs in the Kuroko mine area. It is also intensely and hydrothermally altered. Shikazono et al. (1995) studied the hydrothermal
Miocene-Pliocene Hydrothermal Ore Deposits
i/,
- N-
~11A.8
1"
,' e~
~'~
37
;
/ - 18.8
/
!/
I d 15-4
|
• | lS.7
~11::0
|
" ~ ~' "4
I~:~
I
I
'
~ ZONE I ] IL 13.3• -
~ 10,1 ~"11 6
.O ~ ~= ~ , , f 130
"~
SHINSAWA I°6s 78
•I
! 3.0
014.4
/
i _r~ 133" 11.8
f/ •• ~
[I
19.3 v.~17n .... 13.9
I
18.7
..2~2~. ..... I e14.3 11 ~FUKAZAWA t
~,
es: 3
ZOGAKURA ~l
1~.5 •
",~
!1o.8
,\
KANAYAMAZAWA ~
~ ~=='%1 lS.8 , ~ ~1o.7 TAKARAKURA t/[ III • t~z.8 ~,~ 58
\
...-'~',,~9 -
~4.6 5.6
LEGEND
8,5 010.5
I
~
~~8.8
SARUMA " /
''~
--A'
09.5
ZONE h Zeolite Zone Ih M o n t m o r i l l o n i t e Z o n e IIh S e r i c i t e - C h l o r i t e Z o n e
0 I
z
1 km I
: Kuroko Ore Body : Vein Deposits e6.0 : 8180 .R.(%o) Value
17.8 12.0
10.3 11.9
Figure 1.27. Areal distribution of the whole-rock8180 values of footwall volcanicrocks in the Fukazawa area. The boundaries for the alteration zones are modifiedfrom Date et al. (1983) (Green et al., 1983). alteration of basalt overlying the Kuroko orebody in the Furutobe mine area of Hokuroku district. They showed that hydrothermally altered basalt can be divided into chlorite-rich rock and epidote-rich rock. Chlorite-rich rock occurs widely, whereas epidote-rich rock occupies a smaller area, close to the orebody. It was found that the CaO, Na20 and SiO2 contents of the bulk rock correlate negatively to MgO content, while FeO and NFe contents correlate positively to MgO content (Fig. 1.32) and these changes can be explained by seawater-basalt interaction at elevated temperature. The MgO/FeO ratios of
Chapter 1
38
.
Mont.
Facies
[] Mudstone
Ser.+ Chl.
Basalt, Andesite
Facies
0
[] Tuff J • Dacite
5
10
15
20
25
8180 (%o) Figure 1.28. Whole-rock 3J80 values of Miocene volcanic and sedimentary rocks from the Hokuroku district, grouped by alteration zones. Each square represents one sample. Mont. = montmorillonite, Ser. = sericite, Chl. = chlorite, av. = average (Green et al., 1983).
chlorite and actinolite and the Fe203 content of epidote from the basalt are greater than those of midoceanic ridge basalt (Figs. 1.33 and 1.34) probably owing to the differences in the FezO3/FeO and M g O / F e O ratios of the parent rocks. The lower CaO content and the higher N a 2 0 content of the bulk rock compared with altered midoceanic ridge basalt are interpreted in terms of the difference in original bulk rock composition. 3D of epidote in hydrothermally altered basalt is in a range of -36.5%~ to 43.0%o (Shikazono et al., 1995). Using -37.5%0 as an average 3D value of epidote, the fractionation factor for the H isotope exchange reaction between epidote and H 2 0 (Graham et al., 1980), and 280°C as the temperature of formation of epidote estimated from fluid inclusion study was calculated to be +2.3%~. This value is close to that of hydrothermal fluid issuing from the East Pacific Rise 21°N (+2.0%~, Bowers and Taylor, 1985). This strongly suggests that epidote formed at a discharge zone in submarine hydrothermal system (Shikazono, 1984). It is also notable that the 3D values of epidote from the basalt from Kuroko mine area are close to those of hydrothermally altered midoceanic ridge basalt in the Costa Rica Rift and Galapagos Ridge (-31%~ to -45%~) (Kawahata et al., 1987). 1.3.3. Geochemical characteristics
A large number of geochemical studies on Kuroko deposits (fluid inclusions, gas fugacities, chemical and isotopic compositions of ore fluids etc.) have been carried out. These are summarized below.
Miocene-Pliocene Hydrothermal Ore Deposits 25
[6m0i.÷6-'
20
39
1
'/
'
(at
J
,
% 2~
1
I
4
'l
t
I
~
l
,
to
,
.
.
.
.
,
loo
.
m i 8 0~= 0%°
(b) Zeol. Zone
T=C---2 W e48,,~f 45 o UW.R.
200~, "t !::
5
....
~
.04
.1
2 5
I
'
r '
I
400
t0 '
'
'
,
I
.
.
.
.
(c)
¢:t8n f 15 u ~W.R.
(%.)
"r c-,-
£1
i
I
ii i
Water / Rock ( Atomic Oxygen) Figure 1.29. Calculated changes in the ~180 values of volcanic rocks (~I80 = +7.0%o) as a result of equiIibrium oxygen isotope exchange with waters of different initial compositions. The dotted areas represent the 3180 ranges of rocks in the zeolite and the sericite-chlorite zones (Green et al., 1983).
1.3.3.1. Fluid inclusions Fluid inclusions from Kuroko deposits were studied first by Tokunaga and Honma (1974) who showed that Kuroko deposits formed in a range of 200-260°C for the siliceous
Chapter 1
40
(N)I~3
]] Hole NO
114
180
lance (Kin} 4 .~ration Zone - . . ZeoI
161
I~8 %54~ 5
? ,
~
5~ I~68 1~1 I03
o Ore Zone 0 Mont
~,
{s}
z
Mont
20f
(t}]
(,.ol :(
1
(S)
(NI
Cu
i
I"
:
!
"
(ppm)
'PP" ':I'
'
~'"
'
KeO (wt.%) 4
Zn:f
mt
(ppm) Ioo MgO
(wt%)
CaO
:61
.
:t (wt%)
z
(wt.%)
OI
SF :t
{ppm)
I
0
Figure 1.30. Comparisons of the ~5180values of footwall dacite, alteration zones, and other geochemical halos along the section in the Fukazawa area (Ohmoto et al., 1983).
ore, 100-240°C for barite and 140-170°C for black and yellow ores. Lu (1969) reported homogenization temperatures of the fluid inclusions in quartz from the Uchinotai-Higashi deposit of the Kosaka mine to be 200-250°C. Watanabe (1970) studied fluid inclusions from Ainai deposits and found that the homogenization temperature for the siliceous ores (190-300°C) is higher than that for the black ores (120-290°C). Marutani and Takenouchi (1978)clarified the variations in homogenization temperature and salinity of inclusion fluids in quartz from stockwork siliceous orebodies at the Kosaka mine (Fig. 1.35; Urabe, 1978). They showed that the temperature decreases stratigraphically upwards from stockwork ore zone (280-320°C) to bedded ore zone (260-310°C). Pisutha-Arnond and Ohmoto (1983) carried out fluid inclusion studies of the stockwork siliceous ores from five Kuroko deposits (Kosaka, Fukazawa, Furutobe, Shakanai, and Matsumine) and revealed that black ore minerals (sphalerite, galena, barite) and yellow ore minerals (chalcopyrite, quartz) formed at 200-330°C and 330 4-50°C, respectively, and salinities of the ore fluids remained fairly constant at about 3.5-6 equivalent wt% NaC1. They analyzed fluids extracted from sulfides and quartz; Na ---0.604-0.16 (mol/kg H20), K = 0 . 0 8 ± 0 . 0 5 , Ca = 0.064-0.05, Mg ----0 . 0 1 3 ± 0 . 0 0 8 , C1 ---- 0.82-t-0.32, C (as CO2) -- 0 . 2 0 ± 0 . 1 5 and less than 6 ppm each for Cu, Pb, Zn and Fe.
Miocene-Pliocene Hydrothermal Ore Deposits
41
0 -1
-2
0
o ^~"~
~-3 0 ¢,0
//~///
.g/°--orJ..o" ,, 0///
+o0,% .~" . o
-4
/I
q. 0 /
0
-5 //
-6
~/
Q
0
o
0
-7 I
5
f
t
i
s
I
i
10
,
~
t
I
I
15 a18 0 ( ~ )
I
I
~
I
20
~
i
,
i
I
,
25
Figure 1.31. Bivariable plot of oxygen versus carbon isotopic compositions of carbonates. Solid circle: magnesite; open circle: dolomite; open square: calcite; A: oxygen and carbon isotopic compositions of igneous carbonates; B: oxygen and carbon isotopic compositions of marine carbonates (Shikazono et al., 1995).
Using homogenization temperature and freezing temperature data (Fig. 1.36) and pressure-temperature diagram of N a C 1 - H 2 0 - C O 2 system, minimum seawater depth at the time of Kuroko mineralization is estimated to be 1,000-2,000 m, if boiling of ore fluid did not occur (Fig. 1.37, Pisutha-Arnond and Ohmoto, 1983). This estimated depth is close to that of seawater associated with present-day hydrothermal mineralizations at back-arc basins such as Okinawa Trough (section 3.3). However, Lu (1983) found that the salinity and homogenization temperature of fluid inclusions from the Uchinotai-East ore deposit varies widely and thought that this variation is due to boiling of Kuroko ore fluids. If his argument was correct, the depth could be estimated to be 1,000-1,500 m. Two hypotheses of seaftoor depth at the time of mineralization have been proposed based on foraminiferal data, ca. 3500 m (Guber and Ohmoto, 1978; Guber and Merrill, 1983) and 1500 m (Kitazato, 1979). Considering seafloor depth of present-day ore formation at back-arc basins and fluid inclusion data mentioned above, shallow seaftoor depth hypothesis (Kitazato, 1979) seems more likely. If the pressure-temperature condition of Kuroko ore fluids was close to the boiling curve, the depth could be estimated to be 1,000-1,500 m, which is similar to that for present-day back-arc mineralization such as Okinawa Trough.
1.3.3.2. Gas fugacities Sulfur fugacity (fs2). If electrum is in equilibrium with argentite, the equilibrium constant for the sulfidation reaction (K1-2), 4 Ag (electrum) + 82 = 2 Ag2S (argentite)
(1-2)
where Ag (electrum) is the silver component in electrum, can be expressed as, KI-2 =
a2 I(a 4 ~ Ag2S/I, AgJS2)
(1-3)
Chapter 1
42 30"
O32
o26
10
I0 %o1~ o4
°z7 9
;i 0 • 6
023
~+
09
01
8
~e
01$
028
191 0 21 027
7
t$O 031
+
:ff
~,11
020 03005 08 01+,
O 6 0 (,..)
lO01~&
5
e2~
II 25
0++
7 ~$1
o23
2~ol 7
08
e
3
02 01 •
:+2om6
32
+o %
~3
)26
50 ols°3 12°
0%
2
o28 o~'
200 °It
t
o30 l
I
I
I
I
I
I
2
4
6
8
I0
12
It+
MgO (wt. %)
mgO
16
(wt.%)
o31
~
e2§ o18
e2++
19
022 11
,o%+:,o,,+,.,0,0+,
.,J
ot
v
0
4 e~
4
o28
~2
o=
2
029
0|
%
O29 22 20] 27 o30 I5
o+.
0
Z •
2
~25
+28
~
6
8
+0
MgO
i
t2
I+
16
2
MgO ( w t . % )
(wt. %)
OZa
70
126
I0
ols
032
016 014
65 o~°
o 23 027
8
,j
,..,+
%
0
oJ1
6
,,o ~oo,,",;"
119
11 120~01
U~?
e16
~0
o+2 02
2~i0Of?
35 30
os~s
2+oi
2~2~o 22 +2s 7 6o°
2s..~g ~o o$ Z~eoI o" °~ 011
g o o3
030 012 o13 029
=tl
v
oa
017
Ol
2
4
6 MgO
I
12
( wt.% )
14
16
,
.
2
8
.
.
10
.
12
}&
McjO ( w t . % )
16
Miocene-Pliocene Hydrothermal Ore Deposits
43
A
I
0
! r~
E
D
Z
10
II
~2
)3
I~
15
m6
J7
Fe 2 0 3 Content (wt.%) of Epidote Figure 1.33. Frequency (number of analyses) histogram for Fe203 (wt%) of epidote from the Kuroko basalt. A: epidote coexisting with albite, B: epidote coexisting with chlorite, C: epidote coexisting with pyrite, D: epidote coexisting with hematite and calcite (Shikazono et al., 1995).
where aAg2S and aAg denote the activity of Ag2S in argentite and the activity of A g in electrum, respectively and fs2 is sulfur fugacity. Barton and T o u l m i n ( t 9 6 4 ) have derived a relationship b e t w e e n fs2, temperature, and the A g c o n t e n t of e l e c t r u m in e q u i l i b r i u m with argentite, u s i n g the equation of W h i t e et al. (1957) for the chemical potential of A g in electrum in c o m b i n a t i o n with the equation
Figure 1.32. The relationship between MgO concentration and other major constituents. Solid circle represents the sample of relatively fresh rock which contains original clinopyroxene. A: H20(+) vs. MgO. B: CaO vs. MgO. C: Na20 vs. MgO. D: K20 vs. MgO. E: SiO2 vs. MgO. F: FeO vs. MgO. G: Fe203 vs. MgO. H: EFe vs. MgO. I: A1203 vs. MgO. (Shikazono et al., 1995).
Chapter 1
44 15
A m J
P;H
10
m
t-
<
"6
$ r~
5
E
I . . . . . .
E-I . . . .
I' I
ill
I11 z.J . l . l . l
0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 MgO/FeO (in wt.%) of Chlorite 15
B
>. ¢,t)
<
_c r~
E z
C
5
i
0
•
I¢IMI~IC c l c l c l IClClClC c l C l c l clclcl IclClClCfc c l c l c l ~ 1 , • .I¢IC1¢1~l~ c1¢1cl
0.2 0.4 0.6 0.8
. ,
1.0 1.2 1.4 1.6 1.8 2.0
MgO/FeO (in wt.%) of Chlorite Figure 1.34. Frequency histogram for MgO/FeO ratios (in wt%) of chlorite from the basalt studied (A) and MORB (B). Data sources are: Shikazono and Kawahata (1987), Humphris and Thompson (1978) (M: Mid-Atlantic Ridge) and Kawahata (1984) (C: Costa Rica Rift, Galapagos Spreading Centre). The data on chlorite from MORB are taken from typical metabasalt and not from quartz-chlorite breccia and veins which formed in a hydrothermaI upflow zone (Shikazono et al., 1987).
of Richardson and Jeffes (1952) for the standard free energy change of reaction, 4Ag (native silver) + $2 ----2Ag2 S (argentite)
(1-4)
The equation derived by them can be expressed as, log f s 2 = (1/4.576T) {- 4 1 9 8 0 + 16.52T - 18.296T log NAg
+ (4(1 -- NAg)2[5650 -- 1600(1 -- NAg) -- 1.375T]) }
(1-5)
where NAg and T denote mole fraction of Ag in electrum and absolute temperature (Kelvins), respectively.
Miocene-Pliocene Hydrothermal Ore Deposits
45
Bedded orebody L-40
~e
,
BSO ,
::.,
:t ,
eoe
L-55
,,
YSO
,
BSO
'
YSO
'
|
i•
(
¢
ee eee eeeeeee
I
I
tee e.e oe ceel.¶•
I
1
i!L
Stockwork orebody L-70
BSO
'
S,
• l eeoe e •e~eee
L- 85
.
BSO ,
YSO
'
|
I
I
I
°Oeeee~eee
I|,;
,o,~o*,e
e
I
t
:|
eee oeoe ole, •lee OQII
k-100
BSO ,
,
,
eel•o••
:;:"F:
....
,
,,"
.
YSO
'
I
i ..e
"e g :t eF,:" b eeee te.eyee.
!
200 ° 300 ° Temperature °C
Figure 1.35. Summarized results of homogenization temperature determination in quartz from Uwamuki No. 4 Orebody shown for Kuroko-type (BSO) and Oko-type (YSO), and siliceous ores and for each level (Marutani and Takenouchi, 1978).
Barton and Toulmin (1964) and Barton (1980) have derived the correction to equation (1)-(5) which is necessary due to the solubility o f Au in argentite as a function o f temperature and electrum composition. Based on this equation and NAg, we can place a limit on fs2 and temperature (Fig. 1.38). This application to ore fluids responsible for Kuroko deposits has been done by Sat• (1969), Kajiwara (1970b) and Shikazono and Shimizu (1988a). B o r n i t e - c h a l c o p y r i t e - p y r i t e assemblage also defines i s z - t e m p e r a t u r e region. Combining the FeS content of sphalerite coexisting with b o r n i t e - c h a l c o p y r i t e - p y r i t e , f s z temperature can be determined (Kouda, 1977).
Chapter 1
46 I
g
Oe
•
i
.
.
.
|
I
.
.
.
.
!
P r i m a r y fluid inclusions JU79SK19
S e c o n d a r y fluid inclusions
0
,.-,.
.
6
o Z
5
O" ¢J
o
(b,
4
sw
1I
e
~2
o?
D
i
Low T
b.
0
Secondary Inclusions (16)
s
&
|
.
.
.
.
• o @
@
. . . .
!
150
. . . .
I
200
. . . .
I
250
Temperature
eQ
300
(°C)
Figure 1.36. Homogenization temperature and salinity of inclusion fluids (Pisutha-Arnond and Ohmoto, 1983)•
•
"
I
Pure Wot =~t-
I I
tso-
LIQUID
//,~'/
-
,,,
=
/
loo
- ,sod
/ //
~'
-.-I=o~o
.
,
I
o - . I.'~
c,O~
/
/
Depth of Seowoter
J
/
/
(meters)
. /
,~
/
,.;.'~1 /
VAPOR
'*
A
Fluid I n c l u s i o n F i l l i n g T e m p s .
100
;~00
300
Temperoture ( % )
I
400
350"¢
Figure 1.37. Estimation of minimum depth of seawater at the time of Kuroko mineralization. P-T diagram of NaC1-H20-CO2 is from Drummond (I981) (Ohmoto et al., 1983).
Miocene-Pliocene Hydrothermal Ore Deposits I
I
|
I
t
47 *
i
!
I
,
-5
-10
O4
09 "0'~
-15
0
--I
-20
-25
I
100
0
!
200
300
Temperature °C Figure 1.38. The formative temperature and f s 2 of the black ore from the Shakanai mine. The arrow mark shows an assumed trend of deposition from earlier to later stage (Kajiwara, 1970b). py: pyrite, bn: bornite, cp: chalcopyrite.
Oxygen fugacity (foz). The relationship betweeen
is 2
and
/O 2 can
be derived from the
following reaction. H z S + 1 / 2 0 2 = H 2 0 + 1/2S2
(1-6)
Equilibrium constant for this reaction (K1-6) is expressed as, 1/2 1/2
K16 = (a.2o/
2)/(aH2s/o2
)
(1-7)
where a = activity. Assuming aH20 = 1 and FHzS (activity coefficient of HzS) = 1, we obtain, log fo2 = log fs2 -- 2 logmHzS -- 2log KI-6
(1-8)
where m = molality. Using this relation, we can estimate fo2, mH2S and fs2 at a given temperature. The temperature and f s 2 c a n be estimated from fluid inclusion homogenization temperatures, and chemical compositions of sphalerite and electrum coexisting with argentite and pyrite. H2S concentration of hydrothermat solution is thought to be controlled by chlorite-pyrite assemblage. It is deduced from the H2S concentration of present-day submarine hydrothermal solution which is 1-10 m m o l / k g H 2 0 (Gamo, 1995). Using estimated fs2, temperature, and mH2S, we can place a limit of f Q .
Carbon dioxidefugacity fc02. CO2 fugacity (fc02) of ore fluids is estimated based on CO2 concentration of fluid inclusions analyzed. By using equilibrium constant of the reaction, CO2(g) + H 2 0 ----H2CO3, and assuming aH20 to he unity, fco2 can be estimated.
Chapter 1
48 ~,t,l,~,,l~[[,I,t[,I,,,tl,,~,l
~,~l,t~[ix,,~lt,~,l~Z,l,~,l
Dolomite+K-mica+~.
0
Magnesite+Kaolinite+Quartz <:3
o
2
0
~--2 0 -r
t~
-4
o
_o =
Caleite+Mg-chlorite +K-feldspar
-6 -8
,,,I
0
50
....
I,,,,I,,,,I,,,jll,,z
100
150
200 Temperature( ~ )
Mg-chlorite "6
250
300
licit
0
50
,lllillLl
100
,I;lll,lllwl~w
150
200
Temperature (*C)
250
300
Figure 1.39. Relationship between fco2, aH2co3 and temperature for silicate--carbonate equilibria. Thermochemical data used for the calculations are taken from Helgeson (1969) (Shikazonoet al., 1998). Magnesite, dolomite and calcite occur in hydrothermally altered rocks near Kuroko orebody. The following equations are used to constrain f c Q values of hydrothermal solutions (Shikazono et al., 1998) (Fig. 1.39). 5CaMg(CO3)2 + (dolomite) =
KA13Si3OIo(OH)2 + 3H20 + 3SIO2 (K-mica) (quartz)
MgsA12Si3OIo(OH)8 + KAISi308 + 5CACO3 + 5CO2 (Mg-chlorite) ( K - f e l d s p a r ) (calcite)
(1-9)
5MgCO3 -t- A12Si2Os(OH)4 -t- 7H20 + SiO2 (magnesite) (kaolinite) (quartz) =
MgsA12Si3Olo(OH)8 + 5H2CO3 (Mg-chlorite)
(1-10)
Since temperature of formation of carbonates can be estimated from homogenization temperature of fluid inclusions in carbonates, we can place a limit of CO2 from the above equilibrium relationships. The estimated CO2 range is 1-0.01 mol/kgH20. 1.3.3.3. Chemical compositions o f ore fluids Assuming chemical equilibrium between Kuroko-forming minerals and Kuroko ore fluids, the chemical composition of Kuroko ore fluids can be estimated, using thermochemical data. Calculations for the purposes of estimating the chemistry of Kuroko
Miocene-Pliocene HydrothermalOre Deposits
49
ore fluids have been carried out by several investigators (Sato, 1969; Kajiwara, 1970b; Shikazono, 1976; Ohmoto et al., 1983). For example, if sphalerite is in equilibrium with aqueous solution, the chemical equilibrium for the following reaction can be derived. ZnS = Zn 2+ + S 2-
(1-11)
Assuming that activity of ZnS is unity, solubility product of ZnS (KI-ll) is expressed as, K i-11 = a z n 2 + a s 2 -
( 1-12)
Activity of S 2- is related to total dissolved sulfur concentration (~S) and pH in the region in which H2S is predominant among aqueous sulfur species. S 2-
+ 2 H + = H2S
(1-13)
The equilibrium relation for this reaction is expressed as,
KI_13=aH2s/(a2+as 2 )
(1-14)
Combining this relation with (1-12), and assuming aH2s = NS, we obtain, 2 K 1-14)= (azn2+]ES)/(a2+Kl-13) Kl-ll=(azn2+au2s)/(aH+
(1-15)
In the Kuroko ore fluids with 1 molal C1 concentration predominant dissolved Zn species are thought to be zinc chloro complexes such as ZnC12. For simplicity, assuming that ZnC12 is predominant among dissolved Zn species (~2Zn = mzncl2=aznc12/YZnCl2, where YZnCl2 = activity coefficient of ZnC12 and it is assumed to be 1, and ~2Zn is total dissolved Zn concentration), the relation between azn2+ and 2 Z n can be derived from the following chemical reaction, ZnC12 = Zn 2+ + 2C1-
(1-16)
The equilibrium constant for (1-16) is, Kl-16 = (KI-11KI-13a2H+/~S)(a21 /NZn)
(1-17)
Therefore, we obtain, log NZn = log K - 2pH + 2 log mcl + 2 log YcI
--
log NS
(1-18)
where K = (K1-I1Kl-~3)/K1-16 and Yc1- = activity coefficient of C I - . Thus, NZn can be calculated if the values of K1-11, Kl-13, KI-16, mci-, Yct- and NS are available. NS in ore fluids is generally in a range of 10-2-10 -3 mol/kg H 2 0 based on IES in present-day geothermal waters and fluid inclusion analytical data (Shikazono, 1972a). YcI- is represented as a function of ionic strength and temperature. Ionic strength is related to salinity which can be approximated as C I - concentration. C1- concentration can be estimated from fluid inclusion study. Concentrations of other metal elements can be also estimated based on the procedure similar to that mentioned above.
Chapter 1
50
An example of the calculated results on the chemical compositions of Kuroko ore fluids is given in Tables 1.5 and 1.6. Because of uncertainties of equilibrium constants, NS, pH, temperature, fo2 and other parameters (activity coefficient, ionic strength, activity of water, pressure), the estimated values of concentrations may have uncertainties of-t-1 in logarithmic unit. However, it can be concluded from the thermochemical calculations and fluid inclusion data that the Kuroko ore fluids have the following chemical features. (1) Slightly acidic. (2) C1- concentration is similar to or higher than that of seawater. (3) Mg concentration is very low, compared with that of seawater. (4) The concentrations of base-metal elements and Ba are considerably higher than those of seawater. Several workers have intended to estimate the chemical compositions of Kuroko ore fluids based on the chemical equilibrium model (Sato, 1973; Kajiwara, 1973; Ichikuni, 1975; Shikazono, 1976; Ohmoto et al., 1983) and computer simulation of the changes in mineralogy and chemical composition of hydrothermal solution during seawater-rock interaction. Although the calculated results (Tables 1.5 and 1.6) are different, they all show that the Kuroko ore fluids have the chemical features (1)-(4) mentioned above. As will be discussed later, the experiments (Hajash, 1975; Mottl and Holland, 1978) and theoretical studies on seawater-rock interaction (Wolery, 1978; Reed, 1983) indicate that the Kuroko ore fluids characterized by (1)-(4) above are formed by seawater-crustal rock interaction at elevated temperatures.
TABLE 1.5 Chemical composition of Kuroko ore solution, estimated based on the solubility data of Kuroko forming minerals (Shikazono, 1985c)
pH Cl Na K Ca Mg Fe SO4
H2S Cu Zn Pb Au Ba Sr T (°C) fo2 (atom)
Kuroko ore solution
Seawater
4.5 :t_0.5 33000 ppm 12500 ppm 4000 ppm >500 ppm 1-10 ppm 300 ppm 10-4 m 5 × 10 3 m 5 ppm 20 ppm 20 ppm 10 3 ppm 30 ppm 1-5 ppm 250°C ~ 50°C
7.89 18800 ppm 10760 ppm 399 ppm 412 ppm I294 ppm 2 x 10 3 ppm 2712 ppm
10 - 3 5 ± I a t o m
0.3 x 10 . 3 ppm 3 x 10 3 ppm 0.03 x 10.3 ppm 0.005 x 10 ? ppm 0.33 x 10 3 ppm 7.9 ppm
Miocene-Pliocene Hydrothermal Ore Deposits
51
TABLE 1.6 Concentrations of sulfur and metal species in the Kuroko ore-forming fluids of pH = 4.5 (Ohmoto et al., 1983) Species
ESO 2-
H2S £Fe ECu IBBa li3Zn EPb EAg H4SiO4
Maximum actual, or minimum value
Concentration (Iog molality)
200°C
250°C
300°C
350°C
Minimum Maximum Maximum Actual Actual Max mum Maximum Minimum Maximum Minimum Maximum Minimum Maximum Minimum Actual
-3.5 -2.0 -2.6 -4.3±0.6 -4.6±1.1 -5.7±0.2 -4.1 -3.0 -3.5±0.6 -4.5±0.6 -4.8±0.6 -5.8±0.6 -5.9±0.3 -6.9±0.3 -2.4
-3.5 -2.0 -2.4 -3.3±0.7 -4.4±1.1 -5.4±0.2 -3.9 -4.7 -3.0i0.7 -4.7±0.7 -4.3±0.7 -6.0~0.7 -4.7±0.4 -6.4i0.4 -2.2
-3.5 -2.0 -2.0 -2.4±0.7 -5.2±1.2 -5.2±0.2 -3.5 -4.2 -2.4±0.7 -4.4±0.7 -3.6±0.7 -5.6±0.7 -4.0±0.4 -6.0±0.4 -2.0
-3.5 -1.5 -0.6 -1.6±0.9 -2.6±1.3 -4.1±0.3 -4.1 -5.4 -2.2±0.9 -5.2±0.9 -2.5±0.9 -5.5±0.9 -2.7±0.5 -5.7±0.5 -1.9
1.3.3.4. Stable isotopes 6D and 6180. 3D and 3180 of the Kuroko ore fluids were estimated based on analyses of fluid inclusions, Kuroko-forming minerals and hydrothermal alteration minerals (e.g., Pisutha-Arnond and Ohmoto, 1983). Estimated 8D and 8180 of Kuroko ore fluids are plotted on 3D-3180 diagram (Fig. 1.40).
40
,
,
.
,
,
,
,
o
,
Kuroke Ore-forming Fl@ds
2O
I
SW ! ~z I Periods ii1,111& IV
Periods I%V
0
B
I I I
~D -20 (%o) -40
- ~-.
~
ndary i n_s J
-60
"Magmatic"FLuid] ' ~
-80 -100 -12
I
-1o
-6
!
-4
I
I
-2 0 6180 (%4
|
I
I
i
2
4
6
8
10
Figure 1.40. Summary of 8D and 8180 values of the Kuroko ore-forming fluids and of low-temperature secondary inclusions (Ohmoto et al., 1983). SW: seawater.
Chapter I
52
They are -30%o to -3%~ for 5D and -2%o to +5%o for ~180. The origin of Kuroko ore fluids based on these data is discussed in section 1.3.4. (]34S ofsulj'qd¢$. A large number of ~348 data on sulfides are available. Figure 1.41 shows the summary of sulfur isotopic compositions of sulfides (chalcopyrite, sphalerite, galena, pyrite) (Shikazono, 1987b). Because pyrite exceeds 80% of total amount of sulfide sulfur in the deposit it is reasonable to assume that 534S of pyrite represents ~34S value of sulfides in Kuroko deposits. ~34S values of sulfide minerals from Kuroko deposits vary widely in a range from -6%0 to +9%o. Pyrite with low ~34S value (-6%0 to 0%o) is
<
GALENA
"5
=E Z
1 -6-5-4-3-2
-
0
1
2
3
4
5
6
,<
7
8
9
10
~3~S(O/oo)
CHALCOPYRITE
"6
Z1 -6
-5
-4
-3
-2
-1
-0
1
2
3
4
5
6
7
8
9
10
~34S(% o)
SPHALERITE
"6
=E Z
1
-6-5
4-3-2-1
-0
1
2
3
4
s
s
~
~
1'0
8~4S(O/c~)
;/ 1
7
PYRITE
,-= ICI E-I E-] -6
-5
-4
-3
~ -2
-1
. -0
. 1
. 2
. 3
4
. 5
6
7
8
9
1'0
~3~S(O/c~)
Figure 1.41. Sulfur isotopic compositions of sulfide minerals from Kuroko deposits (Shikazono, 1987b).
53
Miocene-Pliocene Hydrothermal Ore Deposits
s3~s(%,o) -6 1
mudstones end turfs (ht-qtz zone) .(bar-rich zone) --
•
blaok-ores
-5 1
. i.i . . . . . . . . .
o I
+2 i
f
~= _
•--.
'~ I
+4 i
t
+6 I
+8
t
1
t
1
+10
+22
i
I
__-~_-2-=x_. . . . .
-
+ 24 f
i
~,,~ --
eo
•
bar
,,~ ,.~ ,."
r' ~ j
#
/
yellow-ores pyrite-ores veins and disseminated
~ ]
cp
~
,,, I,,PY
Miooene sea
ores
Figure 1.42. Sulfur isotopic variation and the vertical zonation of ores in the Shakanai No. 1 deposit (Kajiwara, 1971).
framboidal type (Komuro and Sasaki, 1985). The origin of framboidal pyrite is not well understood. However, it seems likely that this type of pyrite formed biogenetically and is not of hydrothermal origin. Most ~34S values of hydrothermal pyrite having euhedral morphology range from +2%0 to +7%0. The approximate ranges of variation of ~34S values of other sulfides are: chalcopyrite; +2%0 to +7%0; galena: 0%0 to +4%0; sphalerite: +2%0 to +7%0 (Sasaki, 1974). Sasaki and Kajiwara (1971) estimated average ~34S value of sulfide minerals to be +4.6%0 based on the measurements of mill concentrates from representative Kuroko deposits. In individual deposits, ~34S of sulfides generally increases stratigraphically upwards (Fig. 1.42). (Kajiwara, 1971). Based on the sulfur isotope evidence, Kajiwara (1971) deduced that the ore solutions underwent a progressive cooling and oxidation due to mixing with seawater. Sulfur isotopic data of separated pyrite as the commonest sulfide mineral (Kajiwara, 1971; Kajiwara and Date, 1971) show different ~348 values for the three sub-types of Horikoshi and Shikazono (1978). The ~34S values of pyrite in the C sub-type deposits are higher than the ~34S values of pyrite from the Y and B sub-types. The ~34S values of pyrite from the Y sub-type seem to be slightly higher than those from the B sub-type. Kajiwara and Date (1971) are of a different opinion: the 334S values from the Kosaka district are higher than those in the Hanaoka district, because all sulfur isotopic data from the C sub-type were obtained in the Kosaka district. The sulfur isotopic data on the obtained Uwamuki deposits of the B sub-type in the Hanaoka district indicate systematic decrease in ~348 passing from the yellow ore (+7%0) to the black siliceous ore (+5%o) (Bryndzia et al., 1983). Kajiwara and Date's data (1971) include three ~348 values of pyrite in the Doyashiki deposit of C sub-type in the Hanaoka district. The main Doyashiki
Chapter 1
54 5/+.9
5'
+40
SD
YU
+
AK +35
~.~.-30
TN06
CO
NK
"KS, f t '
•
rN •.~e
I"IZ HT ll.,rr
''I +"H""",
V
+20 1<13 IN
OG
.15 *5
.10
+15
*20
so(sMow) Figure 1.43. Relationship between 3t80 and ~34S values of barite from some Kuroko and vein-type deposits. Abbreviations are: HN, Hirano; NK, Nagaki; AK, Akakura; YU, Yunosawa; SD, Sado; KY, Katsuyama; TN, Teine; OS, Osarizawa; MZ, Mitsuzawa; HT, Hata; OE, Oe; AI, Akaishi; KN, Kohinata; MT, Miyatamata; KZ, Karuizawa; FN, Funauchi; KO, Koyama; IK, Inakuraishi; and OG, Ogoya. "K.S.R." indicates "Karoko sulfate region" (Watanabe and Sakai, 1983).
deposit was, however, mostly mined out when their samples were collected in about 1969. Furthermore, the mine calls the main Doyashiki of the C sub-type and the satellite deposits of Y sub-type together with the Doyashiki deposits (Takahashi and Suga, 1974).
and 3180 of sulfates. 834S and 8180 of sulfates (anhydrite, gypsum, barite) of Kuroko deposits were reported by Sakai et al. (1970) who indicated that 834S and 8180 values of sulfates are very close to but slightly higher than Miocene seawater sulfate value (834S = +20%o to +21%o; 8180 = 0%o). After Sakai's pioneering work, Watanabe and Sakai (1983) and Kusakabe and Chiba (1983) analyzed large amounts of barite, gypsum and anhydrite for 834S and 8180 and confirmed the conclusion drawn by Sakai et al. (1970) (Fig. 1.43). t]34S
1.3.3.5. Radiogenic isotopes Lead isotopes. Sato and Sasaki (1973) concluded on the basis of a remarkable narrow range in lead isotopic composition of Kuroko ores that lead of Kuroko ore came from deep-seated source which originated from subducting pelagic sediments.
Miocene-Pliocene Hydrothermal Ore Deposits o
38.8"
o
/"~--'~Hokuroku ,,4" j ores / ~ .~
central
38.6.
2ospb 2°4Pb 38.4,
55
Honshu.,,/~'.
o
i. ', I o i
•
382"
.p">.%~ j Z
38.0.
...~/
' / I
•
I• o I"iI I /" 0 /" / / , ~ i . / / I- "II /
~
I
[northern /\Honshu
/"
_.
o Kuroko • Epigenetic
t
rowth curve
207pb 15.6- !
2o4pb
~(7 o
15.5-
~
"?.'"~ "--£'o • /
•
/ ~ Honshu
15.4- centrcl| Honshu
18~
'
1~4
'
2o6pb/2O4pb
ld.6
Figure 1.44. Isotopic composition of lead in black ore (open symbols) and in yellow ore (closed symbols) in the Hokuroku district. The isotopic fields for black ore from the Fukazawa, Shakanai, and Kosaka deposits are outlined (Fehn et al., 1983).
Lead isotopic data on Kuroko deposits, vein-type deposits in Honshu and volcanic rocks are summarized and plotted in Fig. 1.44 (Fehn et al., 1983)• Although lead isotopic compositions of Kuroko ores occupy a narrow isotopic range, within a given ore deposit, black ore has a uniform isotopic composition but is significantly higher in radiogenic lead than yellow ore (Fig. 1.44, Table 1.7; Fehn et al. 1983). On the basis of the lead isotopic distribution, Fehn et al. (1983) concluded that a major part of the lead in Kuroko deposits was derived from igneous rocks, probably volcanic rocks with a significant contribution coming from the underlying preNishikurozawa formations and the yellow ore seems to have a greater igneous rock lead component than does the black ore.
Strontium isotopes. Strontium isotopic compositions (87Sr/86Sr) of anhydrite, gypsum and barite from Kuroko deposits are summarized in Fig. 1.45 (Farrell et al., 1978; Honma and Shuto, 1979; Farrell and Holland, 1983; Yoneda et al., 1993; Yoneda and Shirahata, 1995). 87Sr/86Sr values of anhydrite and gypsum are slightly lower than that of seawater, suggesting that most of the strontium was derived from seawater, but a small amount of
Chapter 1
56 TABLE 1.7 Isotopic composition of leads from Kuroko deposits (Fehn et at., 1983) Sample number
Field number
2°6pb/2°4pb
2°Tpb/2°4pb
2°8pb/2°4pb
W 1112132 K- 105 W 1112125
18.489 18.463 18.428
15.603 15.589 15.567
38.678 38.623 38.561 a
JU 78UM64 JU 78UM67
18.458 18.475
15.584 15.594
38.577 b 38.619 b
76-12-2-3 JU 7SUMI25 JU 78UM 129 CF 082713 CF 111163B CF 11 1163Y
18.469 18.464 18.470 18.473 18.491 18.434
15.594 15.5s9 15.594 15.589 15.604 i5.560
38.618 b 38.592 b 38.618 b 38.622 38.665 38.508 c
CF CF CF CF CF CF CF CF CF CF CF CF
18.550 18.580 18.568 18.565 18.566 18.558 18.560 18.575 18.561 18.546 18.542 18.518
15.621 15.622 I5.607 15.617 15.607 15.615 15.595 15.604 15.618 15.596 15.589 15.592
38.686 38.736 38.729 38.659 38.702 38.698 38.665 38.734 38.709 38.637 c 38.613 c 38.620 a
CF7 11282 CF 11286 S 11
18.527 t8.536 18.479
15.621 13.620 15.583
38.707 38.701 38.584 a
S 30
18.527
I5.617
38.624
S 35
18.539
15.598
38.683
MT9
18.451
15.567
38.517
M 8
18.448
15.585
38.378
Kosaka area
Uchinotai East 1 2 3
Uwamuki 2 4 5
Uwamuki 4 6 7 8 9 10 11 Fukazawa area
Tsunokakezawa 12 13 14 15 16 17 18 19 20 2I 22 23 Matsurnine-Sbakanai
11197 34511 l 11922 I 11810 34516 110197 082610B l 11913 111813 I1184 682610Y 111920
area
Shakanai 4 24 25 26
Shakanai 8 27
Shakanai 11 28
Matsumine 29
Matsuki 30
If not stated otherwise all samples were measured on galena from black ore. a Galena taken from yellow ore; b From Sato et al. (1981); c Yellow ore sample.
Miocene-Pliocene Hydrothermal Ore Deposits
57
TABLE 1.7 (continued) Sample number Other
Field number
2°6pb/2°4pb
2°7pb/2°4pb
2°spb/2°4pb
HA 41
18.521
15.587
38.623
CF 11233
18.525
15.594
38.626
FU 17
18.556
15.603
38.662
deposits
Hanawa
31 Omaki
32 Furutobe
33
T
~10
[ ] Gypsum []
c ,< "6
Anhydrite
5
E "1 Z
1
,
.
.m..N.I 0.7070
N N N ~ ~
N N , 0.7080
N N N
0.7090 87Sr/86Sr
Figure 1.45. Variation of the 87Sr/86Sr ratio of anhydrite and gypsum from Kuroko deposits (Farrell and HoIiand, 1983; Shikazono et aI., 1983).
strontium was from igneous rocks. 87Sr/86Sr ratios of barite are in a range from 0.706 to 0.708, suggesting smaller contribution of seawater strontium than anhydrite and gypsum. 87Sr/86Sr of barite from small Kuroko deposits (Iwami in San-in and Minamishiraoi in Hokkaido) are lower (Yoneda, et al., 1993; Honma and Shuto, 1979; Farrell and Holland, 1983) than that from large Kuroko deposits in Hokuroku district. This indicates that a large seawater circulation did not occur in a small Kuroko mine area and that seawater circulation is important for the formation of Kuroko deposits.
Rare earth elements (REE). Analytical results of REE contents of hydrothermally altered volcanic rocks in Kuroko mine area and Kuroko ores are summarized as follows (Shikazono, 1999a) (Fig. 1.46). (1) Positive Eu anomaly is observed for barite, Kuroko ores, ferruginous chert (tetsusekiei), and hydrothermally altered basaltic and dacitic rocks overlying the Kuroko ores. (2) Negative Eu anomaly is observed for hydrothermally altered dacite underlying the Kuroko ores and anhydrite in the dacitic tuff breccia.
Chapter I
58
Hydrothermally altered il dacite and anhydrite underlying the i Kuroko ores E lO0
Barite, Kuroko
i
ore ant
Hydrothermally altere basalt overlying the Kuroko ores
err i
i i
::
::
Ns.s~
i
NS-~S
::
!
i
~
sa2
"-- NS-33
NS-36
NS-39
NS.2
i La Ce
• Nd
i SmEu
i Tb
i i Yb L~
i La Ce
i SntEu
i Tb
i i Yb Lu
I La Ce
i Nd
i SmEu
i Tb
t t Yb Lu
Figure 1.46. REE patterns of the altered volcanogenic rocks and Kuroko ores. Data sources: Shikazono (1999a). (A) Hydrothermally altered dacite and anhydrite underlying the Kuroko ores. (B) Barite, Kuroko ore and ferruginous chert. (C) Hydrothermally altered basalt overlying the Kuroko ores (Shikazono, 1999a).
(3) N e g a t i v e Ce a n o m a l y is o b s e r v e d for h y d r o t h e r m a l l y altered chlorite-rich basalt o v e r l y i n g the K u r o k o ores. (4) N e g a t i v e Ce a n o m a l y and positive Eu a n o m a l y are o b s e r v e d for epidote-rich altered basalt near the orebody. (5) L i g h t rare earth e n r i c h m e n t is distinct and R E E contents are relatively high for the ferruginous chert. The R E E pattern for fresh v o l c a n i c rocks in the Kuroko m i n e area studied by Dud~is et al. (1983) is shown in Fig. 1.47 which shows no negative Ce and no positive Eu a n o m a l i e s and L R E E (Light Rare Earth E l e m e n t ) are not enriched c o m p a r e d with H R E E
1000 Sample / Chondrite
1000~ ~_S a m p l e / C h o n d r i t e
a
Basaltic Rocks
b
~
Acidic Rocks
I lOO
~
,,,..
I0
1 La
!...... _ " : 2 _ _ _ -
10
Ce
Sm Eu
Tb
YbLu
La C e
Pr
Nd
S m E u G d T b D y Ho E r
Yb
Figure 1.47. REE patterns of the fresh volcanic rocks in the Kuroko mine area. Data source: Dud~s et al. (1983). (a) basalt; (b) acidic rocks (Shikazono, I999a).
Miocene-Pliocene Hydrothermal Ore Deposits
59
(Heavy Rare Earth Element). Therefore, it is considered that negative Ce and positive Eu anomalies in hydrothermally altered volcanic rocks, Kuroko ores, and ferruginous chert and LREE enrichment in the Kuroko ores have been caused by hydrothermal alteration and precipitations of minerals from hydrothermal solution responsible for sulfides-sulfate (barite) mineralization. Negative Eu anomaly is found for anhydrite sample and altered dacite underlying the Kuroko ores. One of the possible explanations for this negative anomaly is selective leaching of Eu by circulating hydrothermal solution of seawater origin. Sverjensky (1984) has shown from thermochemical calculations that Eu 2+ is more abundant than Eu 3+ under the reduced environment at high temperatures. Thus, it is considered that Eu is leached more efficiently from the rocks compared with the other rare earth elements. Alderton et al. (1980) have shown that Eu in the granitic rocks in southwest England is depleted due to the sericitization of feldspar-bearing assemblage. Negative Eu anomaly has been reported on highly silicified volcanic rocks around the volcanic rocks around the volcanogenic polymetallic massive sulfide deposit at Que River, Tasmania (Whitford et al., 1988). Thus, it is plausible that negative Eu anomaly of dacitic rocks was caused by the sericitization. This negative Eu anomaly indicates that the alteration minerals in the volcanic rocks underlying the Kuroko ores did not precipitate from ascending hydrothermal solution which interacted at relatively high temperatures (more than 250°C) and reduced condition (Eu2+/Eu 3+ is more than 1) and have positive Eu anomaly. Date et al. (1983) and Green et al. (1983) have shown based on numerous analytical data that Na, Ca and Sr were depleted from footwall dacitic rocks below the Fukazawa Kuroko deposits and this depletion was caused by the addition of K from ascending hydrothermal solution accompanied by the destruction of plagioclase, K-feldspar and volcanic glass and formation of sericite. This evidence supports the above interpretation: The selective leaching of Eu from footwall dacite was caused by the hydrothermal solution. The REE pattern for anhydrite is different from that of seawater, indicating that anhydrite did not precipitate due to the simple heating of seawater that was suggested by Sakai et al. (1970) and Sato (1973). This REE pattern could be explained in terms of the mixing of hydrothermal solution and cold seawater and low degree of seawater/hydrothermal solution mixing ratio (Shikazono et al., 1983). Negative Eu anomaly is also found in the fresh and altered dacitic rocks (Dudfis et al., 1983). Therefore this negative anomaly in anhydrite is also explained in terms of an influence of sericitization of dacite accompanied by the depletion of Eu. Sverjensky (1984) calculated the dependency of Eu2+/Eu 3+ in hydrothermal solution on f Q (oxygen fugacity), pH and temperature. According to his calculations and assuming temperature, pH and fo2 for epidote-stage alteration of basalt and Kuroko ores (Shikazono, 1976), divalent Eu is considered to be dominant in the rocks and hydrothermal solution. Thus, it is reasonable to consider that Eu in the rocks was removed to hydrothermal solution under the relatively reduced condition more easily than the other REE which are all trivalent state in hydrothermal solution. Thus, it is likely that Eu is enriched in epidote-rich altered volcanic rocks. Probably Eu was taken up by the rocks from Eu-enriched hydrothermal solution which was generated by seawater-volcanic rock interaction at relatively low water/rock ratio.
60
Chapter 1
A negative correlation between Mg content and Ca content of hydrothermally altered basalt and dacite from the Kuroko mine area exists. This correlation indicates that Ca in the rocks is removed to fluid by the exchange of Mg in seawater. Eu may behave in the manner similar to Ca during seawater-volcanic rock interaction because of the similarity of their ionic radii. Positive Eu anomaly is observed for hydrothermal solution issuing from the hydrothermal vent on the seawater at East Pacific Rise (Bence, 1983; Michard et al., 1983; Michard and Albarbde, 1986). Guichard et al. (1979) have shown that the continental hydrothermal barites have a positive Eu anomaly, indicating a relatively reduced environment. Graf (1977) has shown that massive sulfide deposits and associated rocks from the Bathurst-Newcastle district, New Brunswick have positive Eu anomalies. These data are compatible with positive Eu anomaly of altered basaltic rocks, ferruginous chert and Kuroko ores in Kuroko mine area having positive Eu anomaly and strongly support that Eu is present as divalent state in hydrothermal solution responsible for the hydrothermal alteration and Kuroko mineralization. Hydrothermally altered basalt in Kuroko mine area can be divided into chloriterich and epidote-rich one (Shikazono et al., 1995). Chlorite formed under the higher water/rock ratio and lower temperatures than epidote (Shikazono, 1984; Shikazono and Kawahata, 1987). At the stage of chlorite formation, a large amount of seawater cycled (seawater/basalt ratio, 40 by mass) (Shikazono et al., 1995). It is considered that Ce4+/Ce 3+ was high during the interaction of relatively low-temperature seawater and rocks at this stage. Ce is present as trivalent in the original rocks. With the proceeding of the reaction of oxidized fluid (modified seawater) with volcanic rocks, Ce is removed to fluid as Ce 4+ which is stable in oxidized fluid as complexes. Removed Ce 4+ may be incorporated into Mn-hydroxides and oxides (e.g., Goldberg et al., 1963; Courtois and Clauer, 1980). It is commonly observed that the weathered rocks exhibit a negative Ce anomaly (Ludden and Thompson, 1979; Menzies et al., 1979). Although these Mn minerals are not found in the Kuroko mine area, a positive Ce anomaly may be caused by the fixation of Ce 4+ in the altered rocks. The negative anomaly is interpreted in terms of the above consideration. On the contrary, the other rare earth elements (except for Eu) in the rocks are present as trivalent. Thus, it is likely that Ce 4+ is more soluble than the other rare earth elements by the reaction of oxidized seawater with rocks. Thus, the negative Ce anomaly of altered basalt could be interpreted in terms of low-temperature interaction of oxidized and relatively unreacted seawater having a negative Ce anomaly caused by the incorporation of Ce into Mn-hydroxides with rocks at late-stage of submarine hydrothermal activity. The REE data, combined with alteration minerals and concentration of major elements in hydrothermally altered rocks, could be used to reconstruct the structure and evolution of a submarine geothermal system accompanied by Kuroko mineralization (Shikazono, 1999a). At the stage of Kuroko mineralization, evolved reacted seawater enriched in Eu, Ca, and Sr formed at low seawater/rock ratio (ca. 1 by mass) and at relatively reduced condition (Eu2+/Eu 3+ greater than 1). Selective leaching of Eu, Ca and Sr occurred from the dacitic rocks underlying the Kuroko ores. The hydrothermal solution enriched
Miocene-Pliocene Hydrothermal Ore Deposits
61
in Ca, Eu and Sr issued from the discharge zone on the seafloor and sulfides and sulfates were precipitated from such hydrothermal solution due to the mixing with cold seawater. Positive Eu anomaly is strong and REE contents are low for the barite-rich ores, whereas positive Eu anomaly is weak, but REE contents are high for the ferruginous chert ore. This difference suggests that hydrothermal signal (low REE and positive Eu anomaly) was strong and was preserved at the stage of the precipitation of sulfides and barite and was weak and not preserved at the ferruginous chert stage. Probably the high REE concentrations of the ferruginous chert is due to the effect of bottom-seawater mixing causing scavenging of REE in cold seawater by Fe- and Mn-hydroxides and oxides forming the ferruginous chert. This view is consistent with the results obtained by Klinkhammer et al. (1983), Ruhlin and Owen (1986) and Olivarez and Owen (1989) who showed the scavenging of REE in seawater by Fe-oxides and Fe-hydroxides near midoceanic ridges. The effect of bottom-water mixing was discussed also for Archean iron formations and Red Sea metalliferous sediments by Barrett et al. (1988). The REE study indicates that the concentrations of REE, particularly Eu and Ce in altered rocks and ore minerals are useful indicators of oxidation state, intensity of discharging hydrothermal solution and evolutionary stage of submarine hydrothermal activity. 1.3.4. Depositional mechanism and origin of ore fluids
1.3.4.1. Depositional mechanism Some mechanisms of anhydrite deposition in Kuroko deposits. Shikazono et al. (1983) considered the depositional mechanism of anhydrite based on the mode of occurrence, texture, Sr content, nature of the contained fluid inclusions and isotopic composition of Sr, S and O in anhydrite together with the mineralogy of the sekko ore, combined with their experimental study on the patitioning of Sr between coexisting anhydrite and aqueous solution. The following is their discussion on the depositional mechanism of anhydrite. Several mechanisms could be invoked to explain the deposition of anhydrite in Kuroko deposits. These include: (1) recrystallization of gypsum and/or anhydrite of evaporite origin; (2) precipitation of anhydrite due to the cooling of hydrothermal solutions; (3) precipitation of anhydrite due to the boiling of hydrothermal solutions; (4) replacement of calcic minerals, such as feldspars, in volcanic rocks or of calcareous foraminifera in mudstones; (5) simple heating of seawater without interaction with country rocks, either above the seawater-sediment interface (Kajiwara, 1971) or beneath the seawater-sediment interface (Farrell et al., 1978; Farrell and Holland, 1983); (6) heating of seawater accompanied by interaction with country rocks; and (7) mixing of ascending hydrothermal solutions with seawater at a site either above the seawater sediment interface (Sato, 1973) or beneath the seawater-sediment interface. Several of these mechanisms can be ruled out based on the geologic environment of the Hokuroku district, because the geologic environment of the Hokuroku area during the Miocene was quite different from that of areas of evaporite formation. The precipitation of anhydrite from hydrothermal solutions has been studied extensively by various workers (e.g., Marshall et al., 1964a,b). The salinity of the inclu-
62
Chapter 1
sion fluids is less than ca. 5 wt% (~1 mol/kg H20) (e.g., Marutani and Takenouchi, 1978; Pisutha-Arnond and Ohmoto, 1983). Simple cooling of hydrothermal solutions is therefore virtually ruled out as a mechanism of anhydrite deposition. This also holds true for the proposal that anhydrite precipitation was due to boiling of the hydrothermal solutions. Studies of fluid inclusions in minerals from Kuroko deposits have yet to produce evidence of boiling in the inclusion fluids (e.g., Marutani and Takenouchi, 1978). Kumita et al. (1980) have investigated the distribution of foraminifera in the M1 and M2 mudstones that overlie the sulfide horizons of the Shakanai mines and have found that the relative abundance of calcareous foraminifera compared to arenaceous foraminifera decreases toward the sulfide orebody. Higher ratios of calcareous foraminifera to arenaceous foraminifera are found in mudstones stratigraphically higher than the thicker sekko body. This evidence suggests that some of the calcium in the anhydrites may have been derived from calcareous foraminifera in the mudstone. However, anhydrite generally occurs in the uppermost part of the tuff and tuff breccia of the T3 unit, which underlies the M2 and M I mudstones. The quantity of anhydrite in the mudstones is very small. This suggests that most of the calcium in the sekko anhydrites was not derived from the calcareous foraminifera distributed in the mudstones. Foraminifera are not found in the T3 tuff unit, and textures showing the replacement of foraminifera by anhydrite have not been observed. If seawater is simply heated, anhydrite begins to precipitate at approximately 110°C if PH20 is equal to Ptotal. The concentration of Ca 2+, Sr 2+ and SO42- in Miocene seawater was probably similar to that of present-day seawater (Graham et al., 1982). The partition coefficient of Sr between anhydrite and aqueous solution has been measured (Shikazono and Holland, 1983) as has the solubility of anhydrite in seawater (Marshall et al., 1964a,b); the variation in the Sr content of anhydrite with temperature during simple heating of seawater can therefore be calculated, using a value of 0.25 for the partition coefficient, Kdsr, experimentally determined by Shikazono and Holland (1983). Figure 1.48 shows that the Sr content of anhydrite which initially precipitated from seawater at about 110°C, should be about 1500 ppm; the Sr content of the anhydrites from Kuroko deposits is between 200 and 2,000 ppm (Fig. I. 19). The low Sr content of most of these anhydrites is difficult to explain in terms of the simple heating of seawater unless the experimentally determined values of KdSr (Shikazono and Holland, 1983) are not applicable to anhydrite deposition in Kuroko deposits. The relationship between the Sr content and the 87Sr/86Sr ratio of a number of anhydrites from Kuroko deposits is shown in Fig. 1.49. Most of the data points fall close to a trend line along which the Sr content of anhydrites increases with decreasing 87Sr/86Sr. The samples from the Fukazawa mines include sekko and paragenetically late vein anhydrites. The line marked A in Fig. 1.49 represents the position of anhydrites precipitated during the heating of Miocene seawater if KdSr had a value of 0.24 during the process of anhydrite precipitation in Kuroko deposits. The large separation between line A and the analytical data for the Kuroko anhydrites is striking. Miocene seawater was apparently not the only source of Sr in these anhydrites. Unfortunately, the processes by which the trend line of the anhydrite compositions was generated are still impossible
Miocene-Pliocene Hydrothermal Ore Deposits
63
3000
"6 E
~2000 c 0
0 1000
w
60
8
I
I
i
I
i
1O0 120 140 Temperature('C )
l
160
Figure 1.48. Change in the strontium content of anhydrite precipitated during the heating of normal seawater without any seawater-rock interaction (Shikazono et al., 1983).
87Sr/86Sr
0.709
R=0.1 0.2
0.3
0.4
0.5. .
. . .
A
B
C /
0.708 I
0.707
[l ./
lJ !a t it
1!
tl
0.706
0.705
lb!
5 0
i
J
1000
2000 S r c o n t e n t (pprn.)
Figure 1.49. Change of the strontium content and 87Sr/86Sr ratio of Kuroko anhydrite during the deposition and dissolution due to the mixing of hot ascending solution and cold solution (normal seawater) (Shikazono et aI., 1983). R mixing ratio (in weight) = S.W./(S.W.+H.S.) in which S.W. and H.S. are seawater and hydrothermal solution, respectively. Open triangle: Fukazawa deposits, Solid triangle: Hanawa deposits, Open square: Wanibuchi deposits, Solid square: Shakanai deposits. Concentration of Ca 2+, Sr 2+ and SO 2 of H.S. are assumed to be 1,000 pprn, 1 ppm, and 10 -4 mol/kg H20, respectively. Concentrations of Ca 2+, Sr 2+ and SO 2 of S.W. are taken to be 412 ppm, 8 ppm, and 2,712 ppm. Temperatures of H.S. and S.W. are assumed to be 350°C and 5°C (Shikazono et aI., 1983).
64
Chapter 1
to define uniquely. The anhydrites containing the lowest concentration of Sr have an isotopic composition close to that of Miocene seawater. This can be explained in three ways: (1) by a value of KdSr much lower than 0.24 during the precipitation of anhydrite in Kuroko deposits, (2) by the reaction of seawater with country rocks, and (3) by the mixing of seawater with one or more solutions in which the Sr/Ca ratio is much smaller than that of seawater. The first alternative is unlikely but not impossible. Past experience suggests that Kd Sr probably depends on the degree of supersaturation of the solutions with respect to calcite during calcite precipitation (Katz et al., 1972). In the experiments by Shikazono and Holland (1983), the solutions from which anhydrite was deposited were considerably supersaturated. It is therefore possible that the values of KdSr extracted from their experimental data are higher than those which controlled the incorporation of Sr in anhydrite during the formation of Kuroko deposits. Experiments at very low degrees of anhydrite supersaturation are needed to determine whether this is a possible explanation for the low Sr content of some of the Kuroko anhydrites. The reaction of seawater with country rocks is also a possible but unlikely explanation. Tertiary volcanic sediments in the vicinity of Kuroko deposits are altered and tend to have lost both Ca and Sr (Farrell and Holland, 1983). The ratio of Sr loss to Ca loss is roughly equal to the Sr/Ca ratio in seawater. If seawater was the altering medium, its St/Ca ratio was probably not strongly affected by the alteration process. The 87Sr/86Sr ratio would be intermediate between an initial value of 0.7088 and ca. 0.740 the 87Sr/86Sr ratio of unaltered Tertiary volcanics of the Hokuroku basin. It is unlikely, therefore, that this type of alteration can account for the Sr content and for the isotopic composition of Sr in the anhydrites at the upper end of the trend line in Fig. 1.49. On the other hand, mixing of seawater with solutions which have a St/Ca ratio much smaller than that of seawater could have led to the deposition of Kuroko anhydrites. Mixing with solutions containing high concentrations of Ca 2+ and very little Sr 2+ or SO] could lead to the precipitation of anhydrite whose Sr content reflects the low St/Ca ratio of the resulting mixtures. Their isotopic composition would be close to that of seawater. If the solutions that mixed with seawater contained low concentrations of Sr with an isotopic composition of ca. 0.7050, anhydrites precipitated from such mixtures would have tended to lie along a family of curves such as curve B in Fig. 1.49. The position of these curves depends on the concentration of Sr 2+, Ca 2+, and SO 2- in the mixing solution as well as on the isotopic composition of its Sr. It seems unlikely that mixing with a series of such solutions is a reasonable explanation for the trend of the data for Kuroko anhydrites in Fig. 1.49. It seems more likely that this trend was generated by the mixing of solutions from which the low Sr, high 87Sr/86Sr anhydrites were deposited with solutions characterized by a considerably higher Sr/Ca ratio and an SVSr/S6Sr ratio of ca. 0.7070. These solutions could have been the hydrothermal solutions from which the barites and presumably the sulfides in the several Kuroko mine areas were deposited. The 87Sr/S6Sr ratio of barites falls in the range 0.7069 to 07079 (Farrell and Holland, 1983): unfortunately the Sr/Ca ratio of the solutions from which the barites were deposited is poorly constrained. If two solutions containing Sr and Ca are mixed, and if no Sr or Ca containing phase is precipitated, the 87Sr/86Sr ratio of the mixtures is related to their -
-
Miocene-Pliocene Hydrothermal Ore Deposits
65
Sr/Ca ratio by the relationship:
{(87 Sr/86Sr) _ (87 Sr/86Sr)1 }/{(87 Sr/86Sr) _ (87 Sr/86Sr)2 }
,~{(mca2+/msr2+)-(mca2+/msr2+)l}/{(mca2+/msr2+)-(mca2+/msr2+)2
} (1-19)
where m denotes molality. The bar indicates the properties of the mixtures, and the subscripts 1 and 2 indicate those of the two end-member solutions. If the amount of anhydrite precipitated during mixing is sufficiently small so that the Sr/Ca ratio of the solutions is not thereby affected, the Sr content of the anhydrites will be related to the isotopic composition of the contained Sr by curves such as curve C in Fig. 1.49. If the quantity of anhydrite precipitated from the mixture is sufficiently large so that the Sr/Ca ratio of the solutions is affected significantly, the Sr content of the anhydrites of any given 87Sr/S6Sr ratio will be greater than those along the calculated mixing curve. The data for the concentration and the isotopic composition of Sr in the analyzed Kuroko anhydrites are generally consistent with such a mixing model. Unfortunately, the model is not unique, and additional chemical and/or isotopic constraints are needed to test its validity. Ohmoto et al. (1983) and Kusakabe and Chiba (1983) also reached the conclusion that the 8180 vs. 87Sr/86Sr relationship and ~34S vs. temperature relationship of barite from the Fukazawa deposit in the Hokuroku district may be explained by a mixing model with a seawater contribution of less than 20% at temperatures around 200°C. Sato (1973) and Ohmoto et al. (1983) calculated the amounts of sulfides precipitated due to the mixing of ascending hydrothermal solution and cold seawater. Their calculations showed that the calculated ratios of the amounts of minerals precipitated are generally consistent with those in Kuroko ore deposits. At early stage of mineralization, anhydrite formed by the mixing of seawater with hydrothermal solution which interacted in shallow part with volcanic rocks at seawaterdominated condition by decreasing Sr/Ca ratio of seawater but not so changing 87Sr/86Sr. In this stage Mg-chlorite formed from such hydrothermal solution dominantly originated from seawater. After this stage, sulfides, barite, and quartz precipitations took place at the sub-seafloor and on the seafloor by the rapid mixing of hydrothermal solution with cold seawater. Rapid precipitation of minerals occurred from the supersaturated solution under non-equilibrium condition. Equilibrium processes cannot explain the following two points. (1) ~34S values of coexisting sulfate and sulfides, and (2) barite/quartz ratio in orebody. If sulfur isotopic equilibrium between coexisting sulfates and sulfides was attained, using average ~34S values of sulfates and sulfides, +22%o and +5%o, respectively, we could estimate temperature using the equation by Ohmoto and Rye (1979). This temperature seems too high compared with temperature estimated from fluid inclusions and mineral assemblages (section 1.3.3). That means that sulfates and sulfides precipitated under the condition far from equilibrium. If hydrothermal solution in which H2S is dominant aqueous sulfur species and is free mixed with cold seawater in which high amounts of SO]- are contained,
SO]-
Chapter 1
66 6O 5O
• Black ore
40
• Yellow ore
30
• Siliceous ore L)
20
¢=
10
•
00
#~ 10
~., 20
$ 30
=-40
, 50
. 60
70
Quartz Content [wt% ]
Figure 1.50. Quartz and barite content of Kurokoore. but no H2S, it is thermochemically predicted that both quartz and barite precipitate with increasing the cold seawater/hydrothermal solution ratio because solubility of quartz decreases with decreasing of temperature and that of barite decreases with increasing S O ] - concentration, which means decreasing of temperature. However, the barite and quartz contents of Kuroko orebody do not positively correlate with each other (Fig. 1.50).
Precipitation of barite and quartz. Barite and quartz are the most common gangue minerals in the submarine hydrothermal ore deposits such as Kuroko deposits and backarc basin deposits (e.g., Okinawa, Mariana deposits) (Halbach et al., 1989; Shikazono, 1994; Shikazono and Kusakabe, 1999). These minerals are also common in midoceanic ridge deposits. The observations of hydrothermal vents at midoceanic ridges and back-arc basins indicated that the minerals precipitate from the hydrothermal solutions which mix with cold ambient seawater. The mineralogical studies of the chimneys from these areas clarified that metastable phases (e.g., amorphous silica, wurtzite, marcasite, native sulfur) are common in the chimneys. The metastable phases are thought to have formed from the highly supersaturated solution. The degree of supersaturation (or saturation index) for barite in Kuroko deposits is estimated to be less than 100 (Shikazono, 1994). The degree of supersaturation with respect to quartz solubility is high because amorphous silica is precipitating from the solution. The above lines of evidence suggest that the precipitation of minerals in submarine hydrothermal ore deposits on the seafloor is taking place from the fluids with high flow rate at the orifices of the chimney (ca. 1-10 m/s) and with high degree of supersaturation under the non-equilibrium conditions. Gamo (1995) revealed based on the chemical and isotopic compositions of hydrothermal fluids from midocean ridges that the precipitation of minerals and interaction
Miocene-Pliocene Hydrothermal Ore Deposits
67
of fluids and sediments under the seafloor affect the chemistry of fluids discharging from the seafloor. Shikazono et al. (1983) indicated that the anhydrite in Kuroko deposits formed at subseafloor environments. Some applications of the coupled fluid flow-reaction model were carried out to the ore-forming process (e.g., Lichtner and Biino, 1992). However, a few attempts to understand quantitatively the precipitations of minerals from flowing supersaturated fluids in the submarine hydrothermal systems have been done (Wells and Ghiorso, 1991). Wells and Ghiorso (1991) discussed the silica behavior in midoceanic ridge hydrothermal system below the seafloor using a coupled fluid flow-reaction model. The behavior of silica and barite precipitation from the hydrothermal solution which mixes with cold seawater above and below the seafloor based on the thermochemical equilibrium model and coupled fluid flow-precipitation kinetics model is described below. As noted already, Kuroko deposits are characterized by the following zonal arrangement in ascending stratigraphic order: siliceous ore (quartz, chalcopyrite, pyrite), yellow ore (chalcopyrite, pyrite), black ore (sphalerite, galena, barite), barite ore (barite and quartz) and ferruginous chert ore (microcrystalline quartz, hematite). Quartz is abundant but barite is poor in the siliceous ore, though barite veinlets occur in this zone. Barite is common in the black ore and abundant in barite ore. Barite is also found in ferruginous chert ore (Kalogeropoulos and Scott, 1983). Quartz is poor in massive sulfide ore horizons (yellow and black ores). The distribution of quartz and barite in Kuroko deposits suggests that barite and quartz precipitate by different mechanisms. Figure 1.50 shows the relationship between barite and quartz contents of the Kuroko ore samples. This suggests that quartz and barite formed separately in different parts (Ohmoto et al., 1983). Amorphous silica and barite precipitate simultaneously from white smoker in midoceanic ridge hydrothermal system (Edmond et al., 1979). It is inferred that amorphous silica precipitates in the chimney at a later stage than sulfides and sulfates (anhydrite and barite) which constitute chimneys from which black smoker is emerging. It is thought that the precipitation of amorphous silica is caused by conductive cooling from the hydrothermat solution which flows laterally in the chimney (Herzig et al., 1988). Barite is abundant in back-arc basin hydrothermal system such as Okinawa, Manus and Mariana (Shikazono and Kusakabe, 1999). In these chimneys, coprecipitation of barite and amorphous silica is taking place from the solution characterized by lower temperatures and lower flow rate than the black smoker. Solubilities of quartz and amorphous silica in aqueous solutions increase with increasing of temperature (Holland and Malinin, 1979). Solubility of barite depends on salinity and temperature (Blount, 1977). The solubility of barite in hydrothermal solution having more than 1 molal NaC1 concentration increases with increasing temperature, while a solubility maximum exists in the solution with NaC1 concentration less than ca. 0.2 molal (Blount, 1977).
Chapter 1
68
==
./
<
om
0
1O0
200
300
400
500
600
Precipitated Amount of Quartz (mg/kg • H20) Figure 1.51. Relationship between precipitated amount of quartz and that of barite.
Salinity of ore fluids responsible for Kuroko deposits is in a range of 0.5-1 tool/1 which is estimated based on the fluid inclusion studies (e.g., Marutani and Takenouchi, 1978). Thus, it seems likely that the precipitation of barite takes place by decreasing of temperature and increasing of concentration of sulfate ion caused by the mixing of hydrothermal solution having 0.5-1 molal NaC1 with cold ambient seawater. Quartz precipitates also due to the mixing of hydrothermal solution with ambient cold seawater if equilibrium between solution and quartz is attained. The thermochemical equilibrium calculations on the amounts of minerals precipitated due to the mixing of ascending hydrothermal solution with cold seawater have also shown that the amounts of quartz precipitated correlate to that of barite precipitated (e.g., Ohmoto et al., 1983; Janecky and Seyfried, 1984; Bowers et al., 1985). Previous studies on the precipitations of barite and quartz calculated the saturation index with respect to barite and quartz during the mixing of hydrothermal solution and seawater. However, the amounts of barite and quartz precipitated were not obtained. Therefore, the precipitated amounts due to the mixing was calculated. The relationship between the amounts of barite and quartz precipitated is shown in Fig. 1.51, which indicates that these amounts are positively correlated. However, as ah'eady noted, the barite content in Kuroko ore inversely correlates to the quartz content and the occurrences of barite and quartz in the submarine hydrothermal ore deposits are different. The discrepancy between the results of thermochemical equilibrium calculations based on the mixing model and the mode of occurrences of barite and quartz in the submarine hydrothermal ore deposits clearly indicate that barite and quartz precipitated from supersaturated solutions under non-equilibrium conditions. Thus, it is considered that the flow rate and precipitation kinetics affect the precipitations of barite and quartz. It was attempted to derive the relationships in the precipitated amounts of barite and quartz, flow rate and precipitation rate using the coupled fluid flow-precipitation
Miocene-Pliocene Hydrothermal Ore Deposits
69
nit J
I
Mixing
Precipitation
/
q C, Figure 1.52. Uniformly mixed kinetics model, q: volume flow rate (m3/s), Ci: initial concentration (molal), l: height (m), r: radius (m), C: concentration (molal), V: volume of the system (m3). model. T w o g e o l o g i c sites o f the precipitations o f barite and quartz are taken into account: one is the site a b o v e the seafloor and the other one is that b e l o w the seafloor. The calculations w e r e m a d e based on the u n i f o r m l y m i x e d precipitation kinetics m o d e l (Fig. 1.52) for one c o m p o n e n t and one d i m e n s i o n a l system r e p r e s e n t e d by Eq. (1-20).
dC / dt = - k ( a / m ) ( c - Co) + (q / V) ( Ci -- C)
(1-20)
w h e r e k = precipitation rate constant, A = surface area on w h i c h barite and quartz precipitate, M = mass o f h y d r o t h e r m a l solution, C = c o n c e n t r a t i o n o f Ba 2+ and H4SiO4 o f the solution in the system, Ci: the c o n c e n t r a t i o n o f input solution, q = v o l u m e flow rate, and V = v o l u m e o f the system. C a l c u l a t e d results are shown in Figs. 1.53 and 1.54. A s s u m e d values o f parameters (l, r, A / M ) for the calculations are g i v e n in Table 1.8. A c o m p a r i s o n o f the calculated results (Figs. 1.53 and 1.54) with the m o d e o f o c c u r r e n c e s o f quartz and barite in the s u b m a r i n e h y d r o t h e r m a l ore deposits indicates
TABLE 1.8 Parameter values used for the computations l (m)
r (m)
A/M (m2/kg)
q/V (l/s)
1
1
1 0.1
0.01 0.1 0.1
0.2 0.02 0.02
v
2 3
v 10 v
v: velocity (m/s); A: surface area (m2); M: mass of fluid (kg); q: volume flow rate (m3/s); V: volume of system (m3).
Chapter i
70
0.16
I
0.12
----
l=lm,r=0.01m l=lm,r=0.1m I I--0.1m,r=-0.1m CBai
O
<
~.~ 0.08
X \ \
0.04
0
\
~"
-,,
~
. . . i . . . .T . . . " r '.- . .---..-t. . . . i,. . . .¢--..
i
2
3
4
5
6
7
8
9
10
Velocity(m/s) Figure 1.53. Relationship between precipitated amount of barite and velocity of fluid at 200°C. v: velocity (m/s), CBai: initial concentration of Ba. 14
12
10
l=Im,r=O.Olm I l=Im,r=O.Im I l=O.Im,r=O.Im Cs o2i
- - - -
81',` "~
4
21| 0 1
-8
\
\",,~
\\\ \ i" - ~
-7
-'-'~...,-,---
-6
- ".~. ~ - -
-5
--~
. . . . .
-4
L------.J
-3
. . . .
-2
-1
Log 1~ Figure 1.54. Relationship between precipitated amount of quartz and velocity of fluid at 200°C. Csio2i: initial concentration of SiO2; v: velocity (m/s).
Miocene-Pliocene Hydrothermal Ore Deposits
71
that the precipitations of barite and quartz take place in different sites in the submarine hydrothermal system. Quartz tends to precipitate under the conditions of high temperature, high A/M, and slow fluid flow rate, while barite under the conditions of low temperature, low A/M, and high fluid flow rate. These results seem to be in agreement with the mode of occurrence of barite and quartz. Barite in the black ore is thought to have formed from the fluids with high flow rate (1-10 m/s) on the seafloor. Barite is found in the chimney from which black smoker is venting. On the other hand, quartz formed in the siliceous ore in high A/M, replacing the dacite below the seafloor. Ferruginous chert in which abundant silica occurs formed below the seafloor by the mixing of ferruginous sediments and hydrothermal components (Kalogeropourous and Scott, 1983). Barite-silica chimney found in back-arc basin formed in the conditions similar to that of ferruginous chert and barite bed in the Kuroko deposits; temperature is relatively low (ca. 150-100°C), and flow rate of fluids may be slow. The above mentioned study on barite and quartz precipitations in Kuroko deposits is summarized as follows. (1) Quartz occurs abundantly in feeder ore (stockwork siliceous ore) in Kuroko deposits. (2) Quartz coexisting with barite also occurs in the ferruginous and barite ores in Kuroko deposits. (3) Barite is abundant in the massive strata-bound ore bodies (black and barite ores) in Kuroko deposits and occurs in the ferruginous chert ore in Kuroko deposits, and chimneys in active deposits at back-arc basins. The coupled fluid flow-precipitation kinetics model calculations indicate the following results: (1) Quartz or amorphous silica tends to precipitate from the solution having relatively high temperature and low flow rate and under high A/M condition. (2) Barite tends to precipitate from the solution with relatively high flow rate and low temperature and under low AIM. (3) These predictions are generally in agreement with the observations; homogenization temperatures of fluid inclusions in quartz from siliceous ore zone and in barite from black ore zone in the Kuroko deposits is relatively high, ranging from 350 to 250°C, and tow, ranging from 250 to 150°C, respectively. (4) The results of calculations are in agreement with the occurrences of barite and silica and chemical features of discharging fluids in the submarine hydrothermal ore deposits; namely, quartz is inferred to precipitate in subseafloor environment and barite in seabottom environment. A/M for the stockwork siliceous ore zone in Kuroko deposits is low, while that for black ore in the Kuroko deposits and chimney may be high. The above-mentioned consideration indicates that important factors controlling the precipitations of barite and silica are surface area/water mass ratio (A/M), temperature, precipitation rate constant (k) and flow rate (v), and the coupled fluid flow-precipitation models are applicable to understanding the distributions of minerals in submarine hydrothermal ore deposits.
Chapter I
72
Barite precipitation highly depends on SO ] . and Ba 2+ concentrations in the fluids. That means that the mixing ratio of hydrothermal solution and seawater is also an important factor for the precipitation of barite, together with the factors mentioned above.
Precipitation mechanism of barite. The precipitation mechanism of barite from aqueous solutions at temperatures below 100°C have been extensively studied by many workers (e.g., Elving and Leineweber, 1950; Wagner and Wuellner, 1952; Nielsen, 1955, 1957, 1958, 1959a,b, 1961; Nielsen and Tort, 1984; Collins and Leineweber, 1956; Takiyama, 1959a,b; Fabrikanos and Lieser, 1962; Lieser and Wertenbach, 1962; Nancollas and Liu, 1975; Klein and Frontal, 1964; Blount, 1974; Liu et al., 1976; Rizkalla, 1983). However, barites in back-arc basin deposits precipitated in the temperatures more than 100°C, such as 150-300°C in Kuroko deposits (e.g., Pisutha-Arnond and Ohmoto, 1983). Thus, Shikazono (1994) carried out the experiments of barite precipitation at elevated temperature (150°C). It was found from his experiments that the morphology of barite varied with the barium chloride concentration. Dendritic crystals with rod-like, spindlelike, and star-like or cross-like habits, and irregularly shaped crystals with rough surface formed from solutions of high barium chloride concentrations (BaCl2, 0.08-0.8 molal). Feather-like dendritic crystal did not form. The results obtained by Shikazono (1994) are generally in agreement with those obtained by previous investigations in their barite precipitation experiments at less than 100°C. For example, Liu et al. (1976) found that dendritic crystals of barite formed at 25°C from aqueous solution containing BaSO4 higher than 2.0 × 10 - 3 molal, which is in agreement with the result by Lieser and Wertenbach (1962). They synthesized feather-like barite from aqueous solutions containing 2.0 x 10.2 molal BaSO4. Suito and Takiyama (1954) and Takiyama (1959a) reported that spindle-shaped dendritic barite crystals having ragged edges formed from aqueous solutions having 1.0 x 10 2 to 2.0 x 10-2 molal BaSO4. Sasaki and Minato (1983) reported that barites precipitated at 60°C from an aqueous solution having the total cation (Ba(NO3)2 + Pb(NO3)2) concentration greater than 0.004 molal were very fine-grained, and complex dendritic particles crystals, while rectangular (Ba,Pb)SO4 crystals formed from a solution of lower concentration. From the values of BaCl2 concentrations and the dissociation constant of BaC1+ (Helgeson and Kirkham, 1976), the concentrations of Ba 2+ in the experimental solutions are estimated to be 0.034-0.047 molal. Assuming the above values of SO42 and Ba 2+ concentrations, the boundary condition for the formation of dendritic barite and wellformed barite is estimated to be - 6 . 2 to 6.0 for the log(mBa2+)i(mso2)i value at 150°C (Fig. 1.55). The log(mBa2+)i(mso2-)i values for the solutions of the previous investigators (T = 25-100 C) and observed morphologies of barite m their experiments) are compared as shown in Fig. 1.55. It shows that the morphology of barite crystals changes with an increase in the concentration product, (mBa2+)i(mso2)i, from well-formed (rectangular, rhombohedra) (R) through rod-hke, star-hke, and spmdle-hke dendrmc (D) to feather-hke dendritic crystals (DF). The concentration products obtained in the experiment by Shikazono (1994) can be compared with the solubility product for large well-formed polyhedral barite crystals o
4
.
•
.
.
.
.
,4
.
.
.
.
.
.
Miocene-Pliocene Hydrothermal Ore Deposits -2
DF
-3
DF
-4
DF
73
D D D D
0-~
D D
m 0"~-8
0 __1
-9 -10 [
I
t
~o
,oo
,5o
i
200 V (°(3)
Figure 1.55. The relationships between the concentration product, (Ba2+)i(SO42-)i,at the initiation of barite precipitation, and morphologies of barite crystals (Shikazono, 1994). The dashed line represents the boundary between dendritic barite crystals and well-formed rhombohedral, rectangular, and polyhedral barite crystals. The 150°C data are from Shikazono (1994); the others from other investigations. D: dendritic (spindle-like, rodlike, star-like, cross-like) barite; DF: feather-like dendritic barite; W: well-formed rectangular, rhombohedral, and polyhedral barite. (i): The boundary between the diffusion-controlled mechanism (Di) and the surface reaction mechanism (S) for barite precipitation at 25°C estimated by Nielsen (1958); ®: The solubility product for barite in 1 molal NaC1 solution at I50°C based on data by Helgeson (1969) and Blount (1977). A-B: The solubility product for barite in 1 molaI NaCI solution from 25 to 150°C based on data by Helgeson (1969).
adopted by Helgeson (1969) and experimentally obtained by Blount (1977). The boundary determined by Shikazono (1994)'s study (log(mBa2+)i(mso2)i = - 6 . 1 ± 0 . 1 at 150°C) 4. lies above the concentration products in 1 molal NaC1 solunon by Helgeson (1969) and Blount (1977) which are plotted as two (the concentration product based on the Blount (1977) experimental results is - 7 . 3 in log unit) for 150°C) and the curve A - B from 25 to 150°C (Fig. 1.55). The above comparison indicates that the solubility of dendritic barite crystals is higher than that of well-formed polyhedral large barite crystals used by Blount (1977). The saturation index for the boundary between dendritic barite and well-formed barite can be estimated to be 10 - 6 ' I / 1 0 .7.3 which is equal to ca. 20. It is usually believed that the growth o f dendritic crystals is controlled by a bulk diffusion-controlled process which is defined as a process controlled by a transportation of solute species by diffusion from the bulk of aqueous solution to the growing crystals (e.g., Strickland-Constable, 1968; Liu et al., 1976). The appearances of feather- and star-like dendritic shapes indicate that the concentrations of pertinent species (e.g., Ba 2+, SO 2 - ) in the solution are highest at the corners of crystals. The rectangular (orthorhombic) crystal forms are generated where the concentrations o f solute species are approximately the same for all surfaces but it cannot be homogeneous when the consumption rate o f solute is faster than the supply rate by diffusion (Nielsen, 1958).
Chapter 1
74
The dependency of rate of precipitation of barite from aqueous solution on time at room temperature studied by Nielsen (1958) suggests that the precipitation of barite from solutions of high levels of supersaturation (i.e., more than 30 as saturation index (S.I.) which is defined as the ratio, (mBa2+)(mso2-)/Ksp, where Ksp is the solubility product m molahty for equlhbrlum) is controlled 15y a bulk diffusion mechanism, while the surface reaction mechanism (polynuclear growth) dominates at low S.I. (i.e., less than 30). The boundary between the bulk diffusion mechanism and the surface reaction mechanism (polynuclear growth mechanism), suggested by Nielsen (1958) at 25°C, is shown in Fig. 1.55 ((~ in the figure). His estimated boundary is consistent with that between dendritic and well-formed barite which was determined by scanning microscopic measurements by other workers (e.g., Fisher and Rhinehammer, 1953; Okada and Magari, 1955; Liu et al., 1976). Nielsen and Toft (1984) summarized the relationship between precipitation mechanism of sparingly solute electrolytes (e.g., BaSO4, AgC1, CaF2) and the degree of supersaturation, distinguishing the surface reaction-controlled and diffusion-controlled regions by plotting growth rates of barite on a PBa-PSO4 diagram (where PBa log EBa; EBa is total dissolved barium concentration; Pso4 log ESO4, a total dissolved sulfate concentration). Nancollas and his group (e.g., Nancollas and Liu, 1975) have suggested second order kinetics for barite precipitation at 25°C with a surface-controlled mechanism even at an S.I. as high as 56. These studies have demonstrated that the surface reaction mechanism dominates to the S.I. of ca. 10-100 and that above this S.I. the bulk diffusion mechanism controls the precipitation of barite. This boundary between two different precipitation mechanisms for barite determined by the experiments at 150°C (Shikazono, 1994) roughly coincides with that between dendritic crystals and well-formed crystals which has been experimentally determined at temperatures lower than 100°C. Several morphologies of dendritic barite, such as feather-like, rod-like, spindlelike, star-like, and cross-like crystals have been recognized when the (mBa2+)i(ms02-)i values were considerably higher than the equilibrium values, although no detailed studies have been made on the relationship between the various morphologies of dendritic barites and the degree of supersaturation. Barite is abundant and widespread in Kuroko deposits. However, it is concentrated especially in the upper horizons (black ore and barite ore), increasing upwards within the black ore. Such a trend is also observed frequently in the tetsusekiei, but rarely in the yellow ore. Barite occurs also as vein-fillings in the stockwork siliceous ore. There are four main types of barite occurrence. Type A is well-formed coarsegrained barite in the black and yellow ores, associated with sulfide minerals (sphalerite, galena, chalcopyrite, pyrite). Fine-grained barite (type B) is also associated with sulfides in the black and yellow ores. Fine-grained barite tends to occur in the upper parts of black ore and also in yellow ore. According to Eldridge et al. (1983), fine-grained sulfides and barite were primitive, rapidly precipitated due to the mixing of ascending hot solution with cold seawater at the seafloor (Sato, 1973; Eldridge et al., 1983). They converted to the coarser-grained barite through dissolution and recrystallization within the ore piles. Fine-grained barite tends to occur in black ore, while coarse-grained barite is found in the .
.
.
.
.
.
4
=
=
- -
- -
Miocene-Pliocene Hydrothermal Ore Deposits
75
lower horizons of black ore and also in yellow ore. Type C is fine-grained barite in the massive barite ore horizon. Under an optical microscope, type C appears as an aggregate of radial crystals, intergrowing with fine-grained quartz. Type C is more elongated in shape compared to types A and B. Most barites (types A, B, and C) are rectangular, lath-shaped, platy, and tabular. Rare type D barite crystals are rectangular and polyhedral. Polyhedral, rhombohedral, or coarse-grained barite crystals (type D) are rarely observed in vugs of black ore or in the inner parts of the chimney recovered from the Hanaoka mine (Shimazaki and Horikoshi, 1990; Shikazono, 1992). The various barite types were examined by scanning electron microscopy (SEM). No dendritic (feather-shaped, spindle-like, rod-like, star-like, or cross-like) barite was found. Coarse-grained barite (type A) and fine-grained barite (type B) have generally rectangular shape but sometimes irregular shapes with rough surfaces. The surface of type B is rougher than that of type than that of type A. Type A and B barites are comprised of very fine-grained, disk-like barite microcrystals. The grain size of an individual microcrystal ranges from ca. 0.1 to 5 Ixm. Massive barite crystals (type C) are also composed of very fine grain-sized (several txm) microcrystals and have rough surfaces. Very fine barite particles are found on outer rims of the Hanaoka Kuroko chimney, while polyhedral well-formed barite is in the inner side of the chimney (type D). Type D barite is rarely observed in black ore. These scanning electron microscopic observations suggest that barite precipitation was controlled by a surface reaction mechanism (probably surface nucleation, but not spiral growth mechanism) rather than by a bulk diffusion mechanism. There are two interpretations for the formation of fine disk-like barite particles which constitute types A, B, and C barites. One is homogeneous nucleation caused by rapid mixing of hydrothermal solution with cold seawater and the coagulation of resulting fine particles. The other is heterogeneous nucleation and growth of barite on the surfaces of pre-existing barite crystal. Calculations of settling velocities for very fine particles in hydrothermal plumes issuing from the ocean floor suggest that very fine (less than 10 g m in diameter) particles cannot settle onto the nearby ocean floor (Shikazono, 1992). However, barite microparticles with grain sizes less than 10 ~ m are common in Kuroko deposits. Therefore, heterogeneous nucleation is likely rather than homogeneous nucleation. Sulfate-sulfide chimneys have been recently discovered from the Mariana back-arc basin (e.g., Kusakabe et al., 1990). They are composed of barite, sphalerite, galena, chalcopyrite, and silica. Barite is most abundant in the chimney. Scanning electron micrographs of barites in a Kuroko chimney show that they are well-formed polyhedra. Morphologically similar barite is also found in the inner side of a chimney from the Mariana back-arc basin. Such barites appear to have formed by recrystallization of fine barites which occur in the outer side of the chimneys. The various morphological features of barites from the Kuroko and Mariana deposits, when combined with the experimental studies on barite precipitation, suggest that the surface reaction mechanism was dominant for the formation of these barites. This implies that the concentration product, (mBa2+)(mso2-), at the initiation of barite precipitation was probably less than ca. 100 times that for equilibrium. This estimate can be evaluated based on the strontium and sulfur isotopic studies
-..-1
TABLE 1.9 Chemical composition of hydrothermal solution experimentally interacted with rocks
T (°C) P (bar} W/R Day's pH Na K Ca Mo Fe Mn Si AI CO2 SO4 H2S Cu Ni Zn Pb Au Hg As Cd Sb Ba Sr
1 (ppm)
2
3
4
5
6
7
8
9
10
11
12
13
14
350
350
455 764 349 3.7 14.9 288 0.04 1475 250 <0.1 0.02 0.3 l 0.17 0.01 0.017 0.02 0.7 0.02 0.04 1.2
1561 1330 2 5.6 8.8 1154 0.05
202 1,5 10.5 8.9 184(I 4.1
4(I(I 1000 5 14 3.7 10750 2300 313 9.0 62.5 54 1275
500 1000 5 14 3.0 10250 460(1 370 5 490 76 1200
500 1000 5 ~4 1.0 11250 950 310 250 7.1 7.1 1500
500 1000 5 14 3.5 8250 1860 1530 14.5 1100 120 1300
500 1000 5 14 3.5 10300 425 1100 7.6 385 70 1100
300 700 3 236 5.35 11084 726 1856 5.6 2.0 2.7 416 <0.26 159 1.39 6.60 <0.05 <0.07 <0.15
400 700 3 ~00 3.85 l 1156 762 1565 10.5 117 53 1684 <(}.26 573 5.65 7.54 <(}.05 <0.07
262 500 50 100 4.65 10300 370 1420 351 45 110 1350 0.21
-
10 16 4.8
300 10(,10 5 t4 3.0 750(I 1150 96 345 25.5 4.6 650
300
10 124 5.2
300 1000 09 16 4.2
2.0 5.5
2.1 6.9
10 40 0.11 0.7(I 1.5 (1.05 0.052 0.06 0.6 0,09 0.08 2.1
1.5
4.1
1 2: graywacke; 3-7: rhyolite: 8: andesite; 9-13: basalt; 14: seawater.
1.1
92 5.6 10963 489 818 0.l 3.5 0,05 580 0.37 118 20 10 0,01 0.01 0.1
0,135
0.50
0.38
<0.15
7.89 10760 399 412 1294 2 x 10 3 0.2 x 10 -3 2900 x 10 3 2 x 10 - 3 145 2717 0.3 x 10 3 0.6 x 10 3 3 x 10 -3 0.03 × 10 -3 0,005 × 10-3 0.1 x 10 3 3 x 10 _3 0.1 x 10 -3 0.33 x 10 -3 20 x 10- 1 7.9 ~"
Miocene-Pliocene Hydrothermal Ore Deposits
77
of barite and strontium content of anhydrite from Kuroko and Mariana deposits by Shikazono et al. (1983), Kusakabe and Chiba (1983), and Kusakabe et al. (1990). These authors indicated that these sulfate minerals precipitated due to the mixing of the hydrothermal solution with cold seawater with a contribution of the end member hydrothermal fluid to seawater being less than 20%. Ohmoto et al. (1983) calculated S.I. of barite to be less than ca. 101'5 during the mixing and cooling of fluids. The estimated value is in agreement with that of Shikazono (1994).
1.3.4.2. Origin of ore fluids Origin of ore fluids is constrained by (1) chemical compositions of ore fluids estimated by thermochemical calculations (section 1.3.2) and by fluid inclusion analyses, (2) isotopic compositions of ore fluids estimated by the analyses of minerals and fluid inclusions (section 1.3.3), (3) seawater-rock interaction experiments, (4) computer calculations on the seawater-rock interaction, and (5) comparison of chemical features of Kuroko ore fluids with those of present-day hydrothermal solutions venting from seafloor (section 2.3). During the last two decades, many experimental studies on the seawater-rock interaction at elevated temperatures (100~00°C) have been conducted. Particularly, detailed seawater-basalt interaction experiments have been done. Several experimental studies on seawater-rhyolite interaction and seawater-sedimentary rock interaction are also available (Bischoff et al., 1981). Examples of chemical compositions of modified seawater experimentally interacted with various kinds of rocks are shown in Table 1.9. Several factors such as C1- concentration, water/rock ratio and temperature are important in controlling the chemical composition of the hydrothermal solution interacted with the rocks. For example, water/rock ratio affects the alteration mineralogy (Mottl and Holland, 1978; Seyfried and Mottl, 1982; Shikazono, 1984). For example, at low water/rock ratio, epidote is stable, while chlorite at high water/rock ratio (Shikazono, 1984; Shikazono and Kawahata, 1987). From rock-water interaction experiments (Table 1.9) and analytical data on fluid inclusions we can derive the relationship between the concentrations of alkali, alkali earth and base-metal elements and concentration of C1- ion in the hydrothermal solutions experimentally interacted with rocks, in the natural hydrothermal solutions and in fluid inclusions (Figs. 1.56-1.58). It is seen in Figs. 1.56-1.58 that the concentrations of elements increase with increasing of C1- concentration. Especially base-metal (Zn, Fe, Mn, Pb) and Ba concentrations increase rapidly with increasing of C1- concentration. This relationship strongly suggests that C1- concentration is a very important factor controlling the chemical compositions of ore solution. Detailed discussion on the relationship between C1- concentration and concentrations of alkali and alkali earth elements were carried out by Shikazono (1978a) (see section 2.1). In Fig. 1.59 the relationship between temperature and concentration of elements (Zn, Ba) at constant C1- concentration which is equal to that of seawater obtained by the experimental studies and analytical data on natural hydrothermal solution (geothermal water) are shown. It is seen that the concentrations of base-metal elements (Zn, Fe, Mn, Cu, Pb) and Ba increase with increasing of temperature. Concentrations of these
Chapter 1
78
Log(Zn) 3 OOX
2 1 x
0
x
x •
x x
x
-;
x x
-3
x
-Z
:;
o
:3
4
,5 Log(CI)
Figure 1.56. Relationship between the zinc and CI- concentration in geothermal waters and hydrothermal solution experimentally interacted with rocks (Shikazono, 1988c).
Loq(Fe)
,"
3
:
," o
,'* ;, , II
G
,~I ' /s " X
1
•
,
st 0
~
o0
,,~f
," X
-2
o0
:2[ -3
0
1
X
~. sesX
sS
-1
,o . ,
!
x, " "
• S
""
•
," e °
oO S.W
:;
:3
4
5
Log(Cl)
Figure 1.57. Relationship between the iron and C] concentration in geothermal waters and hydrothermaI solution experimentally interacted with rocks (Shikazono, 1988c).
Miocene-Pliocene Hydrothermal Ore Deposits
79
Log (Ba) CI = S.W. 4 3 AR. 2
% %
gR.
1
~X
0
X
0
%
-1
% %
% %
-2
% %
% % %
-3
SW
-4
800
1000
600
400
200
T(°C)
0
Figure 1.58. Relationship between Ba2+ concentration and temperature of geothermal waters and hydrothermal solution experimentally interacted with rocks (AR: acidic rocks, BR: basic rock) (Shikazono, 1988c).
Loq(Zn) CI = S.W. 5 4 3 I
2
6
%
1
%
~X
0
% ,%
-1
;-. X"
-2
% %
x
SW
-31000
I
I
800
I
I
600
i
1
400
I
I
200
,
I
o
T(°C)
Figure 1.59. Relationship between zinc concentration and temperature of geothermal waters and hydrothermal solution experimentally interacted with rocks (Shikazono, 1988c).
elements at hydrothermal conditions (200-400°C) are much higher than those of seawater. Therefore, it is evident that the effect of temperature on the concentrations of ore-forming elements (base metals, Ba) is strong. For example, the concentrations of base metal elements in modified seawater interacted with acidic volcanic rocks at 300°C are Fe, 100
Chapter 1
80
TABLE 1.10 Isotopic compositions of Kuroko ore solution (K.O.), seawater (S.W.) and magmatic water (M.W.) (Shikazono, 1978) K.O.
S.W.
M.W.
3348E,o,S ~348I],r,S ~)34Stota1
+21 tO +24%o +2 to +5%o >+2 tO +5%o
+20%o +20%0
0 to +1%,)
~180 ~D 87Sr/86Sr
0 to +5%0 to 30 to -3%0 0.7062-0.7087
0%o 0%o 0.709
+7 to +9%° --48 to --85% 0.704
334Sz,o,S:Sulfur isotopic composition of oxidized sulfur species. ~34S~,r,S: Sulfur isotopic composition of reduced sulfur species. ~348total:Total sulfur isotopic composition of aqueous sulfur species. ppm; Ba, 10 ppm; Zn, 1 ppm; and Cu, 0.5 ppm (Table 1.9). These concentrations are roughly in agreement with those estimated by solubility calculations on Kuroko ore fluids (Tables 1.5 and 1.6). Isotopic compositions of minerals and fluid inclusions can be used to estimate those of Kuroko ore fluids. Estimated isotopic compositions of Kuroko ore fluids are given in Table 1.10. All these data indicate that the isotopic compositions lie between seawater value and igneous value. For instance, 87Sr/86Sr of ore fluids responsible for barite and anhydrite precipitations is 0.7069~).7087, and 0.7082-0.7087, respectively which are between present-day seawater value (0.7091) and igneous value (0.704~3.705). From these data, Shikazono et al. (1983), Farrell and Holland (1983) and Kusakabe and Chiba (1983) thought that barite and anhydrite precipitated by the mixing of hydrothermal solution with low 87Sff86Sr and seawater with high 87Sr/86Sr. ~34S values of sulfides are +2%o to +7%o, which lie between igneous value (0%o (typical value)) and seawater value (+20%o). This suggests that both seawater sulfur and sulfide sulfur of the igneous rocks were incorporated into the Kuroko ore fluids of modified seawater origin. Probably, the sulfide sulfur of Kuroko ore fluids were derived from partial reduction of seawater sulfur and dissolution of pyrite and anhydrite in the country rocks. It is noteworthy that 334S values of sulfides from small B sub-type are smaller than large C sub-type and Y sub-type. This difference could be explained in terms of the extent of seawater sulfate reduction to H2S. The above argument on the calculation of chemical composition of ore fluids, seawater-rock interaction experiments, and isotopic compositions of ore fluids clearly demonstrates that Kuroko ore fluids were generated by seawater-rock interaction at elevated temperatures. The chemistry of present-day hydrothermal solution venting from back-arc basins and midoceanic ridges (sections 2.3 and 2.4) also support this view. There is another opinion on the origin of Kuroko ore fluids. Sawkins (1982) thought that intrusive felsic magmas were the source of the metals and heat in Kuroko hydrothermal systems. He stressed the contributions of magmatic fluid and seawater in
Miocene-Pliocene Hydrothermal Ore Deposits
81
ore fluids from which Kuroko deposits formed based on 3D, 3180 (Hattori and Sakai, 1979; Hattori and Muhlenbachs, 1980) and salinity data on inclusion fluids. His calculated mixing ratio of magmatic fluid to seawater was 1/9. The problem on his magmatic model is that metal content of magmatic fluid mixed with cooler hydrothermal solution of seawater origin is high and metal sulfides tend to precipitate even under the small mixing ratio of magmatic fluid to hydrothermal solution probably at very high temperatures. Another problem of magmatic hypothesis which has been critically thought by Ohmoto et al. (1983) is sulfur behaviour during this type of mixing. High salinity of Kuroko ore fluids does not solely mean magmatic contribution. Instead salinity variation can be reasonably explained by subcritical boiling of fluids of seawater origin. ~180 and 3D data (Fig. 1.40) do not clearly show the mixing of seawater and magmatic water which was favoured by Ishihara and Sasaki (1978). Ohmoto et al. (1983) pointed out that the first problem of the magmatic hypothesis is the large difference in the S concentrations between Kuroko ore fluids and magmatic fluids; fluids derived from acidic magmas at temperatures >750°C are more likely to contain ~ 1 0 -1 to > 1 M ~2S (Burham and Ohmoto, 1980), whereas Kuroko ore fluids appear to have contained 10-3-10 -2 M ZS at 250 4-50°C. The second and most serious problem pointed out by Ohmoto et al. (1983) is that magmatic fluids with +5%e sulfur which was estimated based on 334S values of granitic rocks in Green tuff region by Ishihara and Sasaki (1978) are unlikely to form sulfides of ~+5%o at temperature <300°C under the chemical conditions of the Kuroko ore fluids. Even if we accept that ~D and 3180 values lie between the two end member ~ values, this feature can be also explained by low ~D and high 3180 solution being generated by low seawater/rock ratio condition. Lead isotopic data support this interpretation; namely, these data clearly indicate leaching of lead from the rocks (Fehn et al., 1983). Horikoshi and Shikazono (1978) indicated that 3D of ore fluids for B sub-type which is located at centre of Hokuroku basin is higher, suggesting large contribution of seawater, while 3D of ore fluids of Y sub-type located at the margin of Hokuroku basin is lower, suggeting meteoric water contribution. Ohmoto et al. (1983) thought that the involvement of meteoric water in Kuroko ore fluids is unlikely because they estimated the entire Hokuroku district was under more than 2,500 m of seawater depth (probably ca. 3,500 m and the district was located more than 50 km away from the nearest island based on the studies of foraminifera assemblage in mudstone overlying Kuroko deposits (Guber and Ohmoto, 1978; Guber and Merrill, 1983). However, the seawater depth is inferred to be shallower probably 1,000-2,000 m from the following reasons. One of the reasons is that the seawater depth of hydrothermal ore depositions recently discovered at back-arc basins (Okinawa Trough, Izu-Bonin, Mariana-Trough, Marius Basin, North Fiji Basin, Lau Basin) is mostly 1,000-2,000 m. Oki and Hayasaka (1978) reported an occurrence of recent arenaceous foraminifera in a shallow submarine caldera, which usually lies at a depth of more than 4,000 m on the bottom of the open ocean. The seawater in the caldera has a lower pH than normal seawater due to the inputs of volcanic gas and hydrothermal solution with low pH. Lower pH raises the carbonate compensation depth (CCD) there. Therefore, the estimate of
Chapter 1
82
E v
o)
7-]:
20
19
18
3"~ ~ 17
I ~.~
I 16
15
time
14
(Ma)
Figure 1.60. Variation of subsidence rate for syn-rift basins in the Uetsu district, northeast Honshu (Yamaji, I990). The line of boxes shows the spatially averaged subsidence rate. The rate after 15 Ma is not clear because of uncertainty in paleobathymetry. However, the rate probably decreased to the order of 10-100 m/m.y. If the rate had been of the order of 1 km/m.y, after 15 Ma, the water depth of the inner arc region at 14 Ma would have been much deeper than modem, young, back-arc basins.
seawater depth based on paleobathymetry using CCD as a reference point in the Miocene open ocean by Guber and Ohmoto (1978) might be invalid (Kitazato, 1979; Matoba, 1983; Yam@, 1990). Kitazato's (1979) estimate of 1,000-2,000 m as the seawater depth of Kuroko formation seems more likely. Akimoto and Hasegawa (1989) also estimated to be 1,000-2,500 m. Their estimated seawater depth is similar to that of present-day back-arc basins from which hydrothermal ventings occur (Okinawa Trough, Mariana). Yamaji (1990) studied paleobathymetry of the Northeast Japan Arc at 17-14 Ma (Fig. 1.60) and showed that the arc subsided rapidly from 16 to 15 Ma (Fig. 1.61). According to his paleobathymetry, it is likely that shallow sea (sublittoral ~ 1 5 0 m and upper bathyal 150-500 m) existed near the site of Kuroko formation at 15-16 Ma. Considering the discussion above, it cannot be ruled out that meteoric water was involved in the Kuroko ore fluids. In fact, ~D and 3180 values of fluid inclusions from Iwami Kuroko deposit, west Honshu (San-in district) are considerably lower than those of seawater, indicating an involvement of meteoric water. However, 3D and 3180 values of Nurukawa Kuroko ore fluids (Y sub-type) are plotted between seawater value and igneous (or magmatic) value (Ishiyama et al., 2001), suggesting an involvement of igneous (or magmatic) fluids and no contribution of meteoric water. Here, igneous fluids mean the fluids controlled by igneous rocks, which were generated under a low water/rock ratio.
Miocene-Pliocene Hydrothermal Ore Deposits
®
83
Q •
S
SS
S
/~'Ooo s~ ,~
N = NON-MARINE SUBLITTORAL (0-150 m) o = UPPER BATHYAL (150-500 m) • = M I D D L E T O L O W E R BATHYAL (500-2500 m) S =
,
100 km
Figure 1.61. Paleobathymetry of the NE Japan Are at 17-14 Ma (circled numbers) (Yamaji, 1990). The arc subsided rapidly from 16 to 15 Ma. The entire inner arc region became at middle bathyal depths until 15 Ma, except that Yuri hill remained at shallow marine depths. The hill was submergeduntil 14 Ma (Matoba, 1981). Stippled area, Aosawa basalt (Tsuchiya, 1988). Hatched area, Yuri hill. Note that submarine volcanoes are neglected in this figure because volcaniclasticrocks usually contain few fossils.
The most important conclusion derived from the isotopic studies mentioned above is that isotopic characteristics of Kuroko ore fluids were caused dominantly by seawater-volcanic rock interaction at elevated temperature and by the mixing of seawater with small portions of igneous water or the hydrothermal solution whose chemical and isotopic compositions are controlled by water-rock interaction under the rock-dominated condition and also small proportion of mixing of meteoric water. However, it cannot be decided at present which processes (degree of seawaterrock interaction or mixing ratio of seawater, igneous water and meteoric water) are important for the generation of Kuroko ore fluids solely from the isotopic studies. But experimental and theoretical considerations on seawater-volcanic rocks interaction and origin of hydrothermal solution at midoceanic ridges suggest that Kuroko ore fluids can be produced dominantly by seawater-volcanic rock interaction.
1.4. Epithermal vein-type deposits Epithermal vein-type deposits can be divided into four types based on total metal produced and metal ratio: base-metal type, precious-metal (Au, Ag) type, Sb-type and Hg-
84
Chapter 1
type based on the association of metals. Table 1.2 summarizes estimated total productions of Au and Ag, and Ag/Au. Ag/Au of base-metal type is higher (58 by weight ratio) than that of Au-Ag type (13 by weight ratio). The characteristic features of base-metal and precious-metal types are summarized below and those of Sb- and Hg-type deposits are described in section 1.7.
1.4.1. Geological characteristics 1.4.1.1. Distribution Epithermal Au-Ag vein-type deposits occur widely in the Japanese Islands (e.g., Northeast Hokkaido, Southwest Hokkaido, Northeast Honshu, Sado Island, Izu Peninsula, Kyushu) (Fig. 1.62). The deposits occur in young (late Miocene, and Pliocene), subaerial volcanic regions and in submarine-altered volcanic regions (so-called Green tuff region). The distribution area is located at arc-arc junction (Kubota, 1994) and margin of Green tuff region (Figs. 1.5 and 1.6). These deposits are the principal gold and silver producers in Japan (Table 1.2). Epithermal base-metal vein-type deposits are distributed in the Green tuff region (Southwest Hokkaido, Northeast Honshu) (Fig. 1.62). The distribution area of this type of deposits is nearly same as that of Kuroko deposits. For example, large deposits (Osarizawa Cu-(Au); Ani Cu-Au; Hosokura Pb-Zn deposits) occur in Northeast Honshu, but are more widely distributed in the Green tuff region than Kuroko deposits. The base-metal vein-type deposits in Northeast Japan occur chiefly in Oligocene and early-middle Miocene submarine strata in the members of the Monzen (60-25 Ma), the Daijima (25-15 Ma), and the Nishikurosawa (16-14 Ma), but sometimes they are found in those of the Onnagawa (13-7 Ma) and the Funakawa (7-2.5 Ma) stages of late Miocene (Fig. 1.63) (Nakamura and Hunahashi, 1970). Their distributions are mainly confined to the marginal part of the depressional sedimentary basin which formed during the Nishikurosawa-Onnagawa stage, and are characterized by the presence of rhyolite and dacite or the presence of the so-called Tertiary granite (Nakamura and Hunahashi, 1970). In Southwest Hokkaido, Cu-Pb-Zn veins and Kuroko deposits are distributed in the centre of a volcanic depression zone where submarine rocks predominate, while Au-Ag veins do not (Shikazono and Shimizu, 1988a). Many Cu-Pb-Zn vein-type deposits are hosted by organic sedimentary rocks such as shale and mudstone but almost all Au-Ag deposits occur in altered volcanic rocks. This difference in the host rocks affects the chemical features of ore fluids (fo2, fs2, fco2) (section 1.4.4). 1.4.1.2. Age of mineralization The age of formation of epithermal vein-type deposits can be estimated from K-Ar ages of K-bearing minerals (adularia, sericite) in veins and in hydrothermal alteration zones nearby the veins. A large number of K-Ar age data have been accumulated since the work by Yamaoka and Ueda (1974) who reported K-At age data on adularia from Seigoshi Au-Ag (3.7 Ma) and Takadama Au-Ag deposits (8.4 Ma). Before their publication on the K-Ar ages of these deposits it was generally accepted that epithermal
85
Miocene-Pliocene Hydrothermal Ore Deposits
Sanru II
I '
Hokuryu
.KonomQi
I
0 I00 200km
Q
Todor¢
Yakumo J okok u
Imaiishizaki
Osorizawa
Jlwoto 4o °
Takachi
11nnai iYatanl IHanda Koruizowa
Mizobe
Takeno
Hoshin~
Omidani tAkc
Bojo
\
~1~Seigoshi ..----Yugashimo --Nowaji .Kawozu
35 o
"~ T ~ Toi :ai~1 L ~Sakoshiodomori
1J
Okuchi Hishikari mgano
2
J
~ushikino
lourQ
iwoto I 133 °
I 138°
Figure 1.62. Location of epithermai-type deposits in Japan (Shikazono and Shimizu, 1988a). l: Green tuff and subaeriaI volcanic region of Tertiary/Quaternary ages, 2: Main PaIeozoic/Mesozoic sedimentary terranes, 3: Main metamorphic terranes. TTL: Tanakura tectonic line, ISTL: Itoigawa-Shizuoka tectonic line, MTL: Median tectonic line. Open circle: epithermaI Au-Ag vein-type deposits, solid circle: epithermal base metal vein-type deposits, open triangle: epithermal Au disseminated-type deposits.
vein-type deposits in Tertiary volcanic zone formed in Miocene age, the same as Kuroko deposits because Miocene rocks host the veins. The K - A r age data are summarized in Figs. 1.64 and 1.65. It is obvious in these figures that (1) ages of formation of epithermal vein-type deposits vary widely from 15 to 1 Ma, but are mostly 6-1 Ma, (2) epithermal vein-type deposits have been formed
Chapter I
86
/
Figure 1.63. Distribution of vein-type deposits in a part of the inner belt of Northeast Japan. A: the sedimentary members of the Onnagawa stage, B: the sedimentary members of the Nishikurosawa stage (Nakamura and Hunahashi, 1970).
[] ° _
>,
Au-Ag Vein
[ ] Cu-Pb-Zn Vein
Kuroko
t-
< "5 E Z
15
10
5
1
Age (Ma) Figure 1.64. Ages of formation of Neogenie base-metal vein-type, Au-Ag vein-type and Kuroko deposits, estimated from K-Ar age and paleontologic data (Shikazono, 1987b).
later than Kuroko deposits (15-16 Ma), (3) base-metal veins have been formed earlier than Au-Ag veins (Sawai and Itaya, 1996) (Fig. 1.65), and (4) the frequency of ages of mineralization has three peaks: 10 Ma, 5-6 Ma, and 2-1 Ma.
Miocene-Pliocene Hydrothermal Ore Deposits
87
Au-Ag vein
I
I
I
r
,
Cu-Pb-Zn vein
°+f +I 0L
Mn vein
F I~
Mn strata-bound
z , ,, Barite velnF-- ]
F-'l
Massive barite Pyrite vein
L,
II,
,,
Kuroko
K-At age(Ma) 10
L 1=
Figure 1.65. Histogram of K-Ar ages of hydrothermal ore deposits in the Shakotan-Shikotsu district, Hokkaido (Sawai and Itaya, 1996).
1.4.1.3. Volcanic activity related to mineralization Volcanic rocks associated with this type of mineralization are mostly andesitic rocks. For example, Watanabe (1990b) found that the age of andesitic rocks (flat lava) becomes younger eastward, correlating to the age of epithermal mineralization in Southwest Hokkaido, suggesting a genetic relationship between andesitic volcanism and mineralization. Izawa and Urashima (1989) showed a good correlation of ages of andesitic volcanic rocks and A u - A g deposits in south Kyushu. Their ranges are 4 Ma-1 Ma. However, west of Izu Peninsula, this correlation is unclear and bimodal volcanism (acidic and basaltic activities) seems related to the epithermal A u - A g mineralization.
88
Chapter 1
1.4.1.4. Metal enriched and metal ratios The ore deposits can be classed into two types based on the types of associated metals: Au-Ag rich deposits (Type A) from which Au and Ag are produced as main products, and base metal (Cu, Pb, Zn, Mn, (Sn), (W), (Bi), (Mo), (Sb)) rich deposits (Type B) from which Au and Ag are recovered as byproducts. The deposits are associated with felsic and intermediate volcanic rocks but generally not with felsic plutonic rocks. In Japan Au-Ag deposits associated with granitic rocks (e.g., Au-Ag vein-type deposits in Kitakami) occur commonly. However, these plutonic-type deposits are not described here. Total tonnages of production of Au, Ag and other associated base metal elements were estimated from various records (e.g., Geological Survey of Japan, 1980), as shown in Table 1.2. The Ag/Au ratio of the Type 1-A is lower than that of the Type 1-B from which base metals other than Au and Ag have been produced. In general, the vein-type deposits which have produced large amounts of the base metal elements especially Pb and Zn produce small amounts of Au, but sometimes large amounts of Ag (e.g., Toyoha, Yagumo, Hosokura, Yatani). The vein-type deposits which are rich in Mn have also high Ag/Au (e.g., Ohe-Inakuraishi, Yagumo, Rendaiji). Some vein-type deposits which have produced Cu tend to have large Ag and Au amounts (e.g., Furokura, Osarizawa, Kishu, Ikuno, Tada), but some small deposits which have produced small amounts of Cu have low Ag/Au ratios (e.g., Chitose, Takanosu, and Isobe-Koyama). These deposits, except Chitose from which very small amount of Cu has been produced, are stockwork in form and not typical vein-type deposits. Based on the total production data on individual mines, total production of Au and Ag and Ag/Au ratio from each type of deposits were calculated (Table 1.2). Type 1-A deposits have produced the largest amount of Au and Ag. Total production of Au from the Type 1-B and 2 is small. Ag production from the Type 1-B is lower than but not so different from that of Type 1-A. Average Ag/Au ratio of Type I-A is about 13 and that of Type 1-B is 58.
1.4.2. Mineralogical characteristics 1.4.2.1. Metal zoning Orebody zoning (Park and Macdiarmid, 1963) is observed in Cu-Pb-Zn deposits. For example, in Osarizawa deposit, which is one of the largest Cu-Pb-Zn deposits in Japan, ore metal zoning from deeper to shallower parts is Cu -+ ZnPb --+ AuAg. In the Yatani deposits, Pb-Zn ore occurs in the deeper part, while Au-Ag ore in the shallower part (Figs. 1.66 and 1.67). However, it is uncertain that Au-Ag vein changes continuously to Pb-Zn vein in deeper parts. Orebody zoning is not observed in Au-Ag veins, although the Au/Ag ratio of ore changes considerably with depth. District and regional zonings (Park and Macdiarmid, 1963) are generally not found in Cu-Pb-Zn mine district nor in Au-Ag mine district. 1.4.2.2. Ore minerals Epithermal base-metal vein-type deposits are characterized by the abundant occurrence of sulfides (chalcopyrite, pyrite, sphalerite, galena), and a scarcity of Au-
Miocene-Pliocene Hydrothermal Ore Deposits
89
Au.AgVein
Fault
/" J
$1IJ
/ ~o H 2
%." "'dz.., ,." , ~ ' - -
--'-¢~-'~----
A Bo
wTu o
t
I
~' WH "
500 m
_~..~P K3 ,K2 - ' ~ - . .
!
- ' " '--Esu ~ ~ -
" eo
\
B Figure 1.66. Distribution of Au-Ag veins and Zn-Pb veins of the Yatani deposits. ETI: East-Tengu No. 1, ET2: East-Tengu No. 2, TI: Tengu No. 1, T2: Tengu No. 2, T3: Tengu No. 3, KH: Kanizawa-Honpi, WT: West-Tengu, WT2: West-Tengu No. 2, WH: West-Honpi, H2: Honpi No. 2, KU: Kanizawa Uwabanhi, YH: Yatani-Honpi, W7U: W7-Uwabanhi, E8U: E8-Uwabanhi, E3U: E3-Uwabanhi, N01F: No. 1 fault, N02F: No. 2 Fault (Shikazono and Shimazaki, 1985).
minerals. Ag-minerals (e.g., argentite, pyrargyrite, polybasite) are commonly recognized. Pyrrhotite, magnetite, hematite, and marcasite are occasionally observed. Electron microprobe analyses have revealed that many varieties of Ag mineral occur in the base metal and Au-Ag deposits. The most important features on the occurrences of Au-Ag minerals can be summarized as follows. Au-Ag vein-type deposits (Type l-A): (1) Se minerals such as naumannite, aguilarite, Se-bearing argenitie, and Se-bearing polybasite are found in some deposits (e.g. Sanru, Koryu, Takadama, Kushikino). (2) Au-Ag-Te minerals are rare throughout all of the studied deposits, but sometimes found in some deposits (e.g., hessite from Fuke, Ohkuchi, Arakawa). (3) The common Ag-Au minerals are argentite, As-Se minerals and Ag sulfosalts (polybasite, pyrargyrite, pearceite). (4) Bi, Pb, Zn, and Sn-bearing Ag minerals have not been found. (5) Electrum is abundant, compared with Type 1-B deposits. Native silver is poor in amounts. Base metal vein-type deposits (Type l-B): (1) The common Ag minerals are argentite, and Ag sulfosalts (pyrargyrite, polybasite). Ag sulfosalts are abundant in the late stage of mineralization and argentite occurs in the early stage of mineralization (e.g., Ohe-Inakuraishi, Toyoha).
Chapter 1
90
NO3F ¢,NO2F Ac
~B
, ~/i/~ /
K3
K2 i
?:
i
"/7 •
1
11 :: WH
, /~w7
,/ I
I I
o I
I
200 m 1
Figure 1.67. Vertical cross section along A-B in Fig. 1.66. Abbreviations are the same as those in Fig. 1.66 (Shikazono and Shimazaki, 1985).
(2) Se and Te minerals have not been found in these deposits. (3) Bi, Pb, Zn, and Sn-bearing Ag minerals are rarely found. (4) Abundance of electrum is small, although native silver is abundant in some deposits (e.g., Toyoha, Ikuno). A large number of analytical data on chemical composition of sphalerite are available (Shikazono, 1974a; Watanabe and Soeda, 1981 ). The FeS content of sphalerite from epithermal base-metal vein-type deposits varies widely mostly from 1 to 20 mol% (Fig. 1.68). The FeS content of sphalerite from epithermal Au-Ag vein-type deposits is low, mostly less than 1 mol% (Fig. 1.68). The FeS content of sphalerite from epithermal Se-type is slightly higher than that of epithermal Te-type. The FeS content of sphalerite depends on iron minerals coexisting with sphalerite. The FeS content of sphalerite coexisting with hematite is low (0.5-3.0 mol%), while that of sphalerite coexisting with pyrrhotite is high (mostly 10-20 mol%; Shikazono, 1975). In individual deposits (Toyoha Pb-Zn, Hosokura Pb-Zn deposits), the FeS content of sphalerite coexisting with pyrite varies widely in a range of 1-15 mot%. Mn and Cd contents of sphalerite from epithermal base-metal vein-type deposits are low except Mn content of sphalerite coexisting with alabandite which contains 4.4 wt% Mn (maximum value) (Shikazono, 1975).
91
Miocene-Pliocene Hydrothermal Ore Deposits
Frequency
Frequency Kuroko-typedeposits
EpithermalAu-Ag vein-typedeposits
m 10
15
20 FeS mole %
1"0
1'5
20 FeS mole °/~
Frequency
~rL,~ EpithermalCu-Pb-Zn
•
5
10
15
20 FeS mole%
Figure 1.68. Iron content of sphalerite from Kuroko, epithermaI Au-Ag vein-typeand epithermal base metal vein-typedeposits (Shikazono, 1977a). The Ag content of electrum from epithermal Au-Ag vein-type deposits is mostly in a range of 40-70 atomic% (Fig. 1.69). The Ag content of electrum from the Se-type varies widely (Fig. 1.70). The average Ag atomic% is 50 to 55. The Ag content of electrum from the deposits associated with Te mineralization (e.g., Chitose, Fuke, Takeno) is low, ranging from 26.0 to 40.6 atomic% (Fig. 1.70). Very few data on the chemical composition of electrum from epithermal basemetal vein-type deposits are available. However, it is evident that the Ag content varies widely (Fig. 1.71). The Ag content of electrum from the Osarizawa and Okuyama Cu deposits is low (NAg (Ag atomic%) = 8.6-17.7), while the Ag content of electrum from Pb-Zn-Mn deposits (Toyoha, Oe, Inakuraishi, and Imai-Ishizaki) is high (NAg = 60-80). Motomura (1986) reported that the Ag content of electrum from these deposits is higher than that from epithermal Au-Ag vein-type deposits. The geochemical implication of the Ag content of electrum is discussed in section 1.4.4. Chemical compositions of tetrahedrite-tennantite from epithermal base-metal vein-type deposits are characterized by (1) wide compositional variations, and (2) higher Zn and Sb contents and Ag and lower Fe, As, and Cu contents, compared with Kuroko deposits (Shikazono and Kouda, 1979).
92
Chapter I
100
0 ¢OLL
I
5O
. L• 0
20
40
60
80
100
NAg Figure 1.69. Frequency histogram for the Ag content of electrum from epithermal Au-Ag vein-type deposits in Japan (Shikazono and Shimizu, 1988a).
The differences in Zn/Fe ratio of tetrahedrite-tennantite in epithermal vein-type and Kuroko deposits and that of sphalerite in these deposits can be interpreted in terms of the following exchange reaction: ((Cu, Ag)10Zn2(As,Sb)4S 13)tet "b (FeS)sph = ((Cu, Ag)10Fe2(As,Sb)4S 13)tet Jr- (ZnS)sph (1-21) where (FeS)sph, (ZnS)sph, are FeS and ZnS components in sphaterite, and ((Cu, Ag)]0 Fe2(As, Sb)4Sl3)tet and ((Cu, Ag) ioZn2(As, Sb)4S13)tet components in tetrahedrite-tennantite, respectively.
Miocene-Pliocene Hydrothermal Ore Deposits Te t y p e d e p o s i t s
• n=5
Date Okuyama Kushikino Agawa Chugu Fuke Okuchi Chitose Takeno Sado Kato
93
.,~n=lO • n=2
--~=n=15 = =-n=69 ~n=lO *n=17 -- = •n=6 -- --= n = l • n=l
Se t y p e d e p o s i t s Sakoshi Hishikari Koryu Chitose Sanru Kushikino Takadama Yatani Nebazawa Omidani
, n=61
u
n=20,~ •
L
L n=22 = n=41--
=
• n=8 =
=
= = n=44 i
I
i
I
I
0
I
I
I
50
I
I
NAg
I
100
Figure 1.70. Ag content (atomic fraction of Ag) of electrum from the Te-type arid Se-type deposits (Shikazono et al., 1990). n: number of analyses.
30-
>20o c-" O" LL
10"
0
2"0
40
60
80
100 NAg
Figure 1.71. Frequency histogram for the Ag content of electrum from epithermai base metal vein-type deposits in Japan (Shikazono and Shimizu, 1988a).
If equilibrium is attained for the above reaction, equilibrium constant (K1-21) is expressed as, KI-2I =
(azntet/aFetet)/(aZnspJaFesph)
(1-22)
94
Chapter 1
where aZn~et, aFetet, aZnsph, and aFesph are activities of tetrahedrite-tennantite and activities of ZnS and FeS in sphalerite, respectively. This equation suggests that the Fe content of tetrahedrite-tennantite positively correlates with that of sphalerite at constant temperature and pressure, indicating Fe and Zn contends of tetrabedrite-tennantite are useful to estimate physicochemical parameters (fs2, fo2 etc.) as well as Fe content of sphalerite, although detailed study on thermochemical properties of tennantite-tetrahedrite solid solution is still needed. The chemical compositions of Ag-minerals have been obtained from several A u Ag and base-metal vein-type deposits (Shikazono, 1978b; Ohta, 1991, 1992). For instance, Shikazono (1978b) found that argentite (acanthite) from epithermal A u - A g vein-type deposits (Seigoshi, Ohmidani) contains appreciable amounts of selenium but that from epithermal base-metal vein-type deposits (Toyoha, Ikuno) does not. High selenium contents of galena and Ag sulfosalts (polybasite, pyrargyrite) from epithermal A u - A g vein-type deposits (Kushikino) were reported by Takeuchi and Shikazono (1984) (Table 1.11). The geochemical implications of selenium contents of sulfides for the physicochemical environment of epithermal ore deposition will be discussed in section 1.4.4.
1.4.2.3. Gangue minerals Quartz is the most abundant gangue mineral. It occurs commonly in Au-Ag and P b - Z n deposits but is scarce in Cu deposits. Chalcedonic quartz coexisting with Au-Ag minerals occurs abundantly in A u - A g deposits. Amethyst is generally rare and occurs as a late-stage mineral in Au-Ag and Pb-Zn deposits. Magnetite is common in P b - Z n - M n and Cu deposits but has not been reported in A u - A g deposits. It commonly coexists with other iron minerals such as hematite, pyrite, pyrrhotite, siderite, and chlorite and also occurs in both the main stage of sulfide mineralization and in the late stage of mineralization. The occurrence of hematite is generally more widespread than magnetite in all types of deposits, especially Cu-Au deposits. It tends to occur in late-stage mineralization, generally later than the sulfide mineralization. Inesite is common in A u - A g deposits, especially in the Izu peninsula (Kato, 1928). Generally it is found associated with highgrade gold-silver ores, as, for example, in the Kakehashi vein of the Kawazu mine. It also occurs rarely in Pb-Zn deposits such as Toyoha and Tatsumata. The occurrence of other M n - C a silicates such as johannsenite, bustamite, rhodonite, pyroxmangite, tephroite, and penwithite has been reported from A u - A g and Pb-Zn deposits, but these minerals are not common. They have not been reported from Cu deposits. Hydrated calcium silicate minerals such as xonotlite, truscottite, and gyrolite are rare but have been reported from several Au-Ag deposits. They do not coexist with A u Ag minerals but instead are found with quartz, carbonates, and johannsenite. However, in the Keisen No. 3-2 vein in the Hishikari Au-Ag deposits, a close association of electrum with truscottite, smectite and calcite is observed (Imai and Uto, 2001). Prehnite is found with A u - A g minerals in the A u - A g vein (Kanisawa vein) of the Yatani P b - Z n - A u - A g deposit but is not found in the Pb-Zn vein (Yatani-Honpi vein).
Miocene-Pliocene Hydrothermal Ore Deposits
95
Calcium silicates such as wairakite, epidote, prehnite, laumontite, and stilbite are common in the wall rocks of some A u - A g deposits in the Izu peninsula. Epidote occurs as a gangue mineral coexisting with sulfides and quartz in some Cu deposits, but none of the other above-mentioned Ca and Mn silicates have been reported from these deposits. Laumontite is a common mineral in propylite, which is the host rock for A u - A g deposits. Other zeolites such as mordenite and dachiardite are not generally common, but they are the main gangue minerals associated with A u - A g minerals in the Ohnoyama and Awagano A u - A g deposits. Chlorite is abundant in Cu-Pb-Zn-rich deposits but is scarce in Au-Ag-rich deposits. Fe chlorite is the most common and Fe-Mg chlorite is subordinate (Shirozu, 1969). Almost all of the chlorite is classified as orthochlorite which can be regarded as part of the clinochlore~taphnite solid solution series. In general, chlorite is intimately associated with sulfide minerals such as sphalerite, galena, pyrite, chalcopyrite, and pyrrhotite. A 7 A septechlorite was reported from the Toyoha Pb-Zn deposits (Sawai, 1980). Interstratified chlorite-smectite and vermiculite-saponite are rather common minerals in A u - A g deposits (e.g., Yoneda and Watanabe, 1981), but they have not yet been reported from other deposits. Smectite commonly occurs in A u - A g deposits, and to a lesser extent in Pb-Zn deposits, but only rarely in Cu deposits. The occurrence of beidelite in the Seigoshi and Ohkuchi A u - A g deposits has been described (Nagasawa et al., 1981), and sericite is a common gangue and alteration mineral in C u - P b - Z n rich deposits, where it occurs as a late-stage mineral (Nagasawa et al., 1976). Sericites (Watanabe et al., 1982) from A u - A g deposits occur in the wall rocks as alteration products. In addition, kaolinite commonly occurs as a late-stage mineral in some A u - A g deposits, but not together with the ore minerals. Sometimes kaolinite has been found as an alteration product of adularia after main-stage mineralization. Carbonate minerals occur in almost all the vein-type deposits. In general, they are more abundant than Ca and Mn silicates, but their abundance varies widely with different types of deposits. Large amounts of Mn carbonates (rhodochrosite and manganoan calcite) occur in P b - Z n - M n deposits, moderate amounts in Pb-Zn and Cu deposits, and small amounts in A u - A g deposits. Calcite is abundant in all types of deposits. Siderite is common in Cu deposits, especially in C u - A u deposits, but it is uncommon in A u - A g and Pb-Zn deposits. Siderite coexists with hematite, pyrite, chlorite, and rarely with magnetite; it is considered to have been precipitated after the main stage of sulfide mineralization. Other carbonates such as ankerite, dolomite, and kutnahorite are not common. Ankerite has been reported from several C u - P b - Z n deposits where it is found with sulfides and also with oxides such as magnetite and hematite without sulfides. Dolomite is found in some A u - A g deposits but is rare in other deposit types. Siderite and ankerite coexist with oxides and sulfides, but calcite generally occurs as a late-stage mineral associated with quartz and sericite. Mn carbonates are found with sulfides in P b - Z n - M n deposits such as Ohe and Inakuraishi, but this assemblage is not common. Carbonates from epithermal base-metal and A u - A g vein-type deposits are different. Mn- and Fe-carbonates are common in base-metal vein-type deposits and calcite is abundant in A u - A g vein-type deposits. Shikazono (1973) revealed that the iron content
T A B L E 1.1 l R e p r e s e n t a t i v e analytical data on ore m i n e r a l s from K u s h i k i n o A u - A g d e p o s i t (Takeuchi and S h i k a z o n o , 1984) S a m p l e No.
Period
Cu
Au
Aq
Zn
Fe
Cd
Mn
Sb
As
S
Se
Total
FeS a (tool %)
Electrum 77080212e, L6
IV
-
65.0
33.0
.
76081806a, L6
lI
-
60.5
38.7
.
.
.
.
.
7 7 0 7 2 6 0 3 a , 1.8
lib
-
60.0
38.5
.
.
77072601a, L8
lla
-
63.0
35.0
.
76081810c, L9
lib
-
57.0
42.0
.
.
.
.
.
.
.
.
99.0
76081809f, L9
llb
-
67.0
34.0
.
.
.
.
.
.
.
.
101.0
Za, L 6
IV
2.63
-
69.5 ~
0.09
0.00
-
-
6.49
2,22
11.84
6.36
99.14
77080212b, L6
IV
3.69
-
67.57
0.(X/
0.00
-
-
8.62
0.74
11.49
6.91
99.02
78031601 c, L B
Ila
3.09
-
69.63
0.00
0.00
-
-
9.44
(t.79
14.81
1.66
99.42
77072602f, L8
Ila
2.67
-
67.69
0.01
0.00
-
-
11.24
0.33
14.32
3.53
99,79
77072602i, L8
Ma
3.01
-
66.24
0.00
0.00
-
-
10.60
0.75
14.90
3.75
99.25
77072602d, L8
lla
3,t8
-
69.52
0.(X)
0.00
-
-
10.05
0,93
15.09
1.20
99,97
7 7 0 7 2 9 0 l c , 1.8
I1
2.77
-
68.32
0.11
0.00
-
-
9,98
0.55
12.31
5.57
99.61
7 7 0 7 2 9 0 B c , 1.8
llb
2.48
-
68,15
0.02
0.00
-
-
9.90
0.25
ll.90
6.71
99.41
7 7 0 7 2 7 0 1 a , 1.8
lib
2.68
-
68.91
0.(X)
0.00
-
-
9.65
0.51
12.26
5.92
99.93
7 6 0 8 1 8 0 9 c . L9
IIb
3.80
-
65.78
0.10
0.31
-
-
9.85
0.37
12.04
7.30
99,55
7 6 0 8 2 3 0 3 c , 1.9
IIb
3.84
-
70.83
0.10
0.00
-
-
9.14
0.36
12.86
2.99
100.12
76081809b, L9
IIb
4.12
-
65.09
0.25
0.63
-
-
9.84
0.54
11.50
7.57
99.52
76081806a, L9
IIb
4.13
-
71.36
0.44
0.26
-
-
7.35
0.46
10.01
5.01
99.02
76081806c, L9
Ilb
4.32
-
68.50
0.25
0.00
-
-
9.18
0.44
12.57
4.68
99.94
. . .
.
.
.
. .
. .
. . .
.
. . .
.
98.t) .
99.2
.
98.5
.
98.0
Polybasite
¢5
T A B L E 1.11 ( c o n t i n u e d ) I
Sample No.
Period
Cu
Au
Aq
Zn
Fe
Cd
Mn
Sb
As
S
Se
Total
FeS ~
~.
(tool % ) Tetrahedrite
77080221a, L6
II
23.68
-
20.20
5.24
1.29
-
26.19
1.25
22.30
0.01
77080212e, L6
IV
23.92
-
19.16
4.81
1.71
-
-
27.79
0,02
22.60
0.134
100.05
77072901c,
18
IIb
23.01
-
20.60
4.34
2.02
-
-
26.07
1.16
23.00
I).06
100.26
77072912c, L8
lib
32.42
-
7.14
5,97
1.05
-
-
27.23
1.71
24.00
0.00
99.52
76081809b, L9
IIb
22.99
-
20.90
3.86
2.57
-
26.07
1.23
23.37
0,25
101.24
76081B06b,
1.9
IIb
28.54
-
12.10
6,48
0.57
-
-
27.95
0.73
23.54
0.02
99.93
76081809a,
1.9
IIb
22.02
-
211.39
3,87
2.64
-
-
26.33
1.34
22.88
0.00
100.47
100.16 ~-~
~z Naumanni~
77080212,
1.6
7 6 0 8 2 3 0 3 , 19
IV
-
-
74.17
-
0.64
25.69
100.50
lib
-
-
80.57
-
8.96
7.65
97.18
100.45
~"
Pyrargy~te
1, L d
Ilb
0.00
-
59.98
0.00
0.00
77072602A, LB
lla
0.00
-
59.82
0.00
0.00
77072602B, L8
IIa
0.00
-
60.53
0.00
0.00
_
m
63.84
m
m
65.13
-
-
22.46
0.37
16.45
1.19
-
22.06
0.56
17.08
0.00
99.52
-
-
22.42
0.15
17.12
0,26
100.48
0.76
0.78
0.05
-
-
32.59
-
0.06
0.00
0.10
-
-
33.11
Sphalerite
77072901a,
18
lib
0.10
77072909e, L8
IV
0.00
77072601e, L8
IIa
0.10
62.72
1.56
0.52
0.00
-
-
33.56
76081801a, L9
lib
0.24
62,93
1.38
0.42
0.111
-
-
33.86
76081802b, L9
IIb
0.18
63.65
1.80
0.39
0.00
-
-
7 6 0 8 1 B 10b, L 9
IIb
0.33
63.29
1.98
0.83
0.00
-
76081810c,
lib
1.73
62,14
2.65
0.73
0.01
-
19
r
98.12
1.2
99.40
0.1
-
98.46
2.5
-
98.84
2.0
33.67
-
99.69
2.8
-
33.08
-
99.51
3.0
-
32.40
-
99.66
2.1
a F e S : C o r r e c t e d v a l u e e x c l u d i n g the e f f e c t o f X - r a y s c a t t e r i n g c a u s e d b y c h a l c o p y r i t e .
"--3
98
Chapter 1
(FeCO3) of Mn-carbonates varies widely in a range of 10-3-10 -1 mole fraction. Siderite from Au-Ag vein-type deposits (Ohmori) contains appreciable amounts of zinc (0.8-5.8 wt% as ZnO) (Shikazono, 1977b). Adularia is abundant in Au-Ag deposits, where it is commonly found with AuAg minerals; only rarely does it occur in Pb-Zn and Cu deposits. Albite is very rare and is reported only from the Nebazawa Au-Ag deposits. Barite is a common gangue constituent in P b - Z n - M n deposits, especially those in the southwestern part of Hokkaido and the northern part of Honshu, where it is usually a late-stage mineral coexisting with carbonate and quartz but rarely with sulfide minerals. Other rare gangue minerals include fluorite, apatite, gypsum, bementite, rutile, and sphene, but they have not been studied. Main gangue minerals of the Se-type deposits comprise quartz, adularia, illite/ smectite interstratified mixed layer clay mineral, chlorite/smectite interstratified mixed layer clay mineral, smectite, calcite, Mn-carbonates, manganoan calcite, rhodochrosite, Mn-silicates (inesite, johannsenite) and Ca-silicates (xonotlite, truscottite). In comparison, the Te-type deposits contain fine-grained quartz, chalcedonic quartz, sericite, barite, adularia, chlorite/smectite interstratified mixed layer clay mineral and rarely anatase. Carbonates and Mn-minerals are very poor in the Te-type deposits and they do not coexist with Te-minerals. Carbonates are abundant and barite is absent in the Se-type deposits. The grain size of quartz in the Te-type deposits is very fine, while large quartz crystals are common in the Se-type deposits although they formed in a late stage and do not coexist with Au-Ag minerals. Principal gangue minerals in base-metal vein-type deposits are quartz, chlorite, Mn-carbonates, calcite, siderite and sericite (Shikazono, 1985b). Barite is sometimes found. K-feldspar, Mn-silicates, interstratified mixed layer clay minerals (chlorite/smectite, sericite/smectite) are absent. Vuggy, comb, cockade, banding and brecciated textures are commonly observed in these veins. The predominant gangue minerals vary with different types of ore deposits; quartz, chalcedonic quartz, adularia, calcite, smectite, interstratified mica/smectite, interstratified chlorite/smectite, sericite, zeolites and kaolinite in Au-Ag rich deposits; chlorite, quartz, sericite, carbonates (calcite, rhodochrosite, siderite), and rare magnetite in Pb-Zn rich deposits; chlorite, sericite, siderite, hematite, magnetite and rare epidote in Cu-rich deposits (Sudo, 1954; Nagasawa et al., 1976; Shikazono, 1985b).
1.4.2.4. Hydrothermal alteration zoning Among the epithermal vein-type deposits in Japan, four major types of hydrothermal alteration can be discriminated. They are: (1) propylitic alteration, (2) potassic alteration, (3) intermediate argillic alteration, and (4) advanced argillic alteration. The definitions of these types of alteration are mainly based on Meyer and Hemley (1967) and Rose and Burt (1979) who classified the hydrothermal alteration in terms of alteration mineral assemblages. Representative propylitic alteration minerals include epidote, albite, carbonates, quartz, chlorite, sericite, and smectite. The less common minerals are mixed-layer clay minerals such as chlorite/smectite and sericite/smectite and zeolite minerals. The term "propylite" is widely used to describe altered volcanic rocks recognized
Miocene-Pliocene Hydrothermal Ore Deposits
99
in mine areas. However, the term is somehow ambiguous, because it is defined differently by different investigators (e.g., von Richthofen, 1886; Kato, 1928). Propylite is defined here as the andesitic, dacitic and rhyolitic rocks which have undergone the alteration to form the characteristic minerals such as chlorite, epidote, albite, K-feldspar, etc. Probably there are two different origins of this type of alteration. The propylite may be formed either by metamophism or by hydrothermal alteration. However, it is extremely difficult to distinguish rigidly between these two origins. Here the studies on the propylitic alteration which is considered to be intimately related to igneous activity in the mine area are summarized. Usually lateral and vertical zonations are observed in the area of propylitic alteration. Generally epidote, actinolite and chlorite tend to occur in the central and deeper parts, while at the marginal and shallower parts zeolites, mixed-layer clay minerals and smectite are commonly observed. The following zonation is generally recognized from the central to marginal parts: chlorite --+ chlorite/smectite -+ smectite. This zoning pattern is observed in the Fuke Au-Ag district (Inome et al., 1981), Seigoshi Au-Ag district (Shikazono, 1985a), Hosokura Pb-Zn district (Suzuki et al., 1982), Toyoha PbZn district (Yoshitani, 1971; Sawai, 1984), Miyatamata, Arakawa and Nissho C u - P b Zn districts (cited in Utada, 1980). At the central parts of some mine districts listed above, intrusive plutonic bodies have been observed. The formations of high temperature alteration minerals including epidote, actinolite, prehnite and wairakite as suggested by Yoshitani (1971), Fujii (1976), and Shikazono (1985a) are considered to be attributed to these intrusive activities. Zeolite minerals (wairakite, laumontite etc.), mixed-layer clay minerals and smecite occur in the upper part of the propylitically altered rocks (e.g., Seigoshi, Fuke, Kushikino), but they are sometimes poor in amounts. Generally carbonates are more abundant in the mine area as in the Toyoha district. Temporal relationship between the formation of high temperature propylitic alteration minerals (epidote, actinolite, prehnite) and low temperature propylitic alteration minerals) (wairakite, laumontite, chlorite/smectite, smectite) in these areas (Seigoshi, Fuke, Kushikino) is uncertain. Potassic alteration has been recognized in the Numanoue (Otagaki, 1951), Konomai (Urashima, 1953), Takarakura (Utada, 1980), Seigoshi (Shikazono and Aoki, 1981; Shikazono, 1985a; Imai, 1986), Hosokura (Suzuki et al., 1982; Konno et al., 1984), Chitose (Takatori and Nohno, 1985), Karuizaiwa (Sugaki et al., 1986) and the Ikutahara district, Hokkaido (e.g., Yahagi, Ryuo) (Matsueda et al., 1992) where adularia occurs dominantly. In the Mikawa (Sudo et al., 1953, Nagasawa, 1961, 1962), Ohe (Tsukada and Uno, 1980), Toyoha (Okabe and Bamba, 1976), Rendaiji (Watanabe and Nagai, 1986) and Kushikino (Izawa et al., 1987), sericitization is dominant. The area of the potassic alteration is not wide, compared with the propylitically altered area. The width of potassic alteration zone away from the vein is generally within several tens of meters (ca. 50 m) (Shikazono and Aoki, 1981; Imai, 1986). The potassic alteration is usually found in the intermediate vicinity of the vein in the epithermal deposits in Japan. Thus it is evident that this type of alteration occurs genetically related to the ore deposition. Lateral zonation from a sericitic envelope to an intermediate argillic envelope is common in the porphyry copper deposits and vein-type deposits in granodioritic rocks
100
Chapter I
(Meyer and Hemley, 1967). However, such lateral and concentric zonation has not been reported from the epithermal vein-type deposits in Japan. Montmorillonite-rich and silicarich zones exist in the upper part of the Au-Ag veins such as the Seigoshi and Takadama (Nagasawa et al., 1981). In contrast to the hardly investigated lateral zonation around Japanese epithermal vein-type deposits, a few examples of vertical zonation are known. Potassic alteration grades upwards into intermediate argillic alteration in the wall rocks for the Toyoha (Okabe and Bamba, 1976), Ohe (Tsukada and Uno, 1980), Chitose (Hasegawa et al., 1981) and Kushikino (Imai, 1986). Advanced arigillic alteration is found at the upper horizon than the sites of potassic and intermediate argillic alterations where the Au-Ag mineralization occurs (e.g., Seigoshi, Yatani, Kushikino, Hishikari). This type of alteration takes blanketform in upper part and vein-form in lower part (Iwao, 1962; Shikazono, 1985a). The conspicuous zonation from upper to lower horizon is known at the Ugusu silica deposit, namely, silica zone, alunite zone, kaolinite zone and montmorillonite zone (Iwao, 1949, 1958, 1962). It is rather difficult to determine the sequence of each type of alteration in a mine area. However, it is widely accepted that the hydrothermal alteration proceeds as follows: propylitic alteration ~ potassic alteration and intermediate argillic alteration -~ advanced argillic alteration. The actual sequence alteration might be more complicated and superimposition of each type of alteration could be common. Usually propylitic alteration precedes the base metal and Au-Ag mineralizations. Potassic and intermediate argillic alterations are nearly contemporaneous with ore deposition. Advanced argillic alteration is also nearly contemporaneous with base-metal and Au-Ag mineralization as found in the Seigoshi and Yatani Au-Ag mine areas based on K-Ar datings (Shikazono, 1985e), although advanced argillic alteration was not caused by the hydrothermal solution responsible for the Au-Ag mineralization. Advanced argillic alteration at the Ugusu silica deposit is inferred to be caused by the mixing of volcanic gas containing SO2 and groundwater (Shikazono, 1985a). Generally, the chemical composition of rocks does not considerably change during the propylitic alteration. The components which are added to the rocks are only H20, CO2 and S (e.g., Okabe and Bamba, 1976). Considerable changes in the chemical composition of rocks occur during the advanced argillic alteration. For example, SiO2 content of highly silicified rocks of the Ugusu silica mine reaches 99% (Iwao, 1962). This silicification is caused by the considerable leaching of elements from the rocks by acid hydrothermal solution except Si and addition of Si from hydrothermal solution.
1.4.2.5. Spatial and geochemical relationships between propylitic alteration and advanced argillic alteration: a case study on the Seigoshi-Ugusu district, central Japan In the Izu Peninsula, located in the middle part of Honshu, more than 20 epithermal Au-Ag vein-type deposits have been mined. Large Au-Ag mines are located in the western part of the peninsula. The Seigoshi mine is the largest one. The country
Miocene-Pliocene Hydrothermal Ore Deposits
101
TABLE 1.12 Generalized stratigraphic succession in the Izu Peninsula and the area surveyed (Shikazono, 1985a)
~ ~e" O° om R" °
' Daruma
andesite lava +500m
iOdoi
andesite lava +500m
~
Tanaba andesite lava .~ =~" Andes te pyroclastics +500m
~'
e- O
K0shimoda andesite lava Andesite volcanic breccia +500m
I-- v
Nekko Dacite
~' "~ ~. R"
dacite lava tuff breccia
+500m I
e= !~
~)
Yagisawa dacite lava andesitic tuff, Formation tuff breccia +1000m •-dacitic tuff u~ breccia andesite lava Tom (~ andesitic tuff, breccia +1000m •= O. Formation sandstone J~ ~ dacitic tuff ~ volcanic conglomerate Ugusu andesite lava pyroclastics >" Formation sandstone conglomerate +500m
rocks in the Seigoshi district have suffered intense propylitic and advanced argillic alterations. The Izu Peninsula is mainly composed of pyroclastic and volcanic rocks of Tertiary-Quaternary age. The general geology of the peninsula has been well studied (Tayama and Niino, 1931), and thus, it is briefly described below. The generalized stratigraphic succession in the Izu Peninsula is shown in Table 1.12. The oldest rock exposed in the peninsula is the Yugashima Group of Miocene age that is composed of submarine andesitic-dacitic pyroclastic and volcanic rocks with small amounts of sandstone. This rock is metamorphosed from zeolite facies to epidoteprehnite-pumpellyite facies. The total thickness of the group is variable from place to place, but, generally exceeds 1500 m. The Shirahama Group which is mainly composed of felsic tuff and sandstone overlies the Yugashima Group conformably or unconformably in different places. The age of this group is considered to be late Miocene to Pliocene. The thickness of this group is also variable from place to place, but the average thickness is more than 1,000 m. These Tertiary rocks are overlain unconformably by subaerial Quaternary andesitic rocks. The generalized geology and schematic stratigraphic succession in the SeigoshiUgusu district is shown in Fig. 1.72. The andesitic and dacitic rocks of Pliocene age unconformably overlie these Tertiary rocks. The Ugusu Formation, corresponding to the lower horizon of the Yugashima group, is characterized by the predominance of andesitic pyroclastic rocks. The maximum thickness of the Ugusu Formation is estimated
Chapter 1
102
l~
....
v v v vVVV'V'V'V
yvVvVvV~vVvVvVvVvVv3yvVvV~vVvVvVvVv vVv ~,Vv v VvVv~v~vV~Vv~v~,,v, v yvV~ v ~WW~VW4 v v v yvV~VvVvV,,Vv~ VvVW,
...........
~AAAA
A A A A A A
A A A A A
~l~
o
0
o o
0 o
o o
0 o
o"o° o
o
0 o
o o
o o
o o
A A ~
Daruma-Odoi Volcano (andesite lava) 1 ~ Tanaba Volcano (andesite lava, pyroclastics) Koshimoda Volcano (andesite lava, volcanic breccia) Nekko Volcano (dacite lava, tuff breccia) Yagisawa formation ........ (dacitic lava, andesitic tuff, tuff breccia dacitic tuff breccia)
[~
Toi formation (andesite lava, andesitic tuff breccia, sandstone, dacitic tuff, volcanic conglomerate) ~ - ~ Ugusu formation (andesite lava, pyroclastics, sandstone, conglomerate) Intrusive rock of diorite porphyry
Figure 1.72. Generalized geological map of the Seigoshi-Ugusuarea (Shikazono, 1985a).
to be more than 500 m. The Toi Formation, corresponding to the upper horizon of the Yugashima Group, conformably overlies the Ugusu Formation. This formation is mainly composed of pyroclastic and volcanic rocks of andesitic composition. The average thickness of this formation is 1,000 m. The Yagisawa Formation, which may correspond to the Shirahama Group, unconformably overlies the Toi Formation. This formation is composed of dacitic and andesitic lavas and pyroclastic rocks. Its total thickness is ca. 1,000 m. Andesitic and dacitic rocks of Pliocene or Pleistocene unconformably overlie the Yugashima and Yagisawa formations. In the northern part of this district, Quaternary andesitic rocks are distributed. These rocks are composed of an alternation of thick lava flows of hypersthene-augite andesite and their autobrecciated parts. Many faults trending generally from north to south are developed, mainly in the Yugashima Group (Fig. 1.73). The axis of the minor foldings in the Yugashima rocks is also generally north to south. Basic and felsic intrusive and dike rocks often occur in the
Miocene-Pliocene Hydrothermal Ore Deposits
t
103
J
,,' ,'~... Advanced Argillic ~" . \ ,' p.. ,' Alteration m T "1 ,' ~Funabara T \ ,'..----0 . . . . 4. area " ~ " : " ~ o ~Seigosi~i'-. " , ~ o i ,• T o"',l ~ ~, ~ "',. . ,,mine . co ~ n e ', ,.~ ,'/ . . , . , Propylltlc 'j .." .,' ', 'J" ,,', ",<~ Alteration .. ,; ~ F9 ,,'~ ,, ~,-" :/Ugusu ,' Mochikoshi /," ,: m ne : \ '. mine
', F / .,a-x" - 2.'. 't
, ;
Amagi mine : , , ' ~ A ' " ?~ ,'
"J'.
" 'J
",3
',.
~, < 3 ,
F
"~
.... -..2~ ', "
5 km
;',. . . . . . . . . . . . ', ~ ' ,
•t
'
'." "
['~2q Diorite-Porphyry ~
~
Silicified Rock of Advanced Argillic Alteration Gold-SilverQuartz Vein Fault
Figure 1.73. Distribution of epithermaI Au-Ag vein-typedeposits, propylitic and advanced argillic alterations and intrusive rocks of diorite prophyry(Shikazono, 1985a).
Yugashima, trending from north to south and from east to west. Occasionally the A u - A g quartz veins are hosted by these intrusive rocks and the upper horizon of the Yugashima Group. These intrusive rocks have suffered propylitic alteration. The host rocks for the advanced argillic alteration are generally younger formations such as Pliocene andesite, although the Yugashima Group rocks have also suffered the advanced alteration occurring as a vein in form. The distribution of the A u - A g vein-type deposits in this district is shown in Fig. 1.73. The propylitic alteration is intimately associoated with these deposits. The vein is composed of rhythmic banding of quartz layers and fine-grained sulfides such as argentite, acanthite, sphalerite, galena, pyrite and chalcopyrite, and electrum. The principal gangue minerals are quartz, calcite, adularia and interstratified chlorite/smectite. Minor minerals are inesite, johansenite, xonotlite and sericite. These gangue minerals except for quartz, adularia, calcite and sericite are not found in the wall rocks. The Seigoshi and Toi deposits occur in the andesitic pyroclastic rocks of the upper horizon of the Yugashima Group and basic intrusive rocks. Distributions of the wallrock alteration minerals from underground in the Seigoshi mine and on the surface near the
Chapter 1
104 West <
> East
',, • Ao~-'. . . . .
...... 2 . . . .
...... \--~ii-Ne~-2-_:-_:-2-_-2yuo
_. . . . ~ , 2 - 1 ..-+
5--4 j
jJ
', ~~ \ "
'
-.
,
,
.:~J/-~,," ~,, opx J
A~
J'~:f-
f~f J'~
•
+-
?I,
o~
\
"
\
'
', ',
I~.
'
/
"
/
T
I
I
I t
e/
',
--
, /l
', ,
#5-, "v ,
.......
~
I
SeigoshiAu-AgVein
Figure 1.74. Zonal sequence of the propylitic alteration in E-W section of the Seigoshi-Toi mine area (Yug = yugawaralite; Heu = heulandite; Stil = stilbite; Opx = orthopyroxene; Mont = montmorillonite;Mor = mordenite; Lm = laumontite;Wr = wairakite; Chl = chlorite; pr = prehnite;ep = epidote; Py = pyrite; Kf = K-feldspar; Cpx = clinopyroxene)(Shikazono, 1985a). Deep(A)
)
Shallow(B)
Mordenite Yugawaralite Heulandite Stilbite Laumontite Wairakite Montmorillonite Chlorite Epidote Prehnite K-feldspar Pyrite
Pyrrhotite Magnetite Sphene
.....
Figure 1.75. Zonal sequenceof the propylitic alteration in section A-B in Fig. 1.74 (Shikazono, 1985a).
mine are shown in Fig. 1.74. This type of alteration has a vertical zonal arrangement (Fig. 1.75). The abundance of the alteration minerals also changes from the portion near the A u - A g - q u a r t z vein to the portion away from the A u - A g - q u a r t z vein. From the deep to shallow and from the portion near the A u - A g - q u a r t z vein to the portion away from the vein, the hydrothermal alteration zoning is observed: (1) epidote-prehnite-K-feldsparchlorite zone; (2) wairakite-laumontite zone; and (3) stilbite-heulandite-smectite zone. The rocks of the underground levels of the Seigoshi mine have suffered alterations (1) and (2). The boundary between these zones cuts the strata with a small angle. The boundary between each zone is gradual. In zone (1), quartz, K-feldspar, epidote, chlorite, prehnite and sphene are predominant alteration minerals. Epidote, prehnite and carbonate replace plagioclase phenocryst. Epidote often occurs as a veinlet with several millimeters wide, together with prehnite. K-feldspar, calcite and quartz tend to occur as a veinlet. Chlorite replaces pyroxene
Miocene-Pliocene Hydrothermal Ore Deposits
105
and also occurs as a veinlet often with pyrite. Orthopyroxene is completely replaced by chlorite, but clinopyroxene is sometimes preserved. Amphibole is often replaced by carbonate. The rim of the original magnetite is replaced by pyrite and sphene. In zone (2), wairakite and laumontite are commonly found. Wairakite occurs as a veinlet together with laumontite. Laumontite occurs as a veinlet and also as filling amygdule. Very small amounts of epidote are found with these zeolite minerals. Small amounts of interstratified sericite/montmorillonite are found. Veinlets of epidote, prehnite, quartz, K-feldspar and chlorite appear at the deeper part of this zone. Yugawaralite occurring as a veinlet is rarely found. In zone (3), stilbite, heulandite, and minor amounts of chabazite and mordenite are found mainly as veinlets and filling amygdule. The rocks of this zone have not been significantly altered. Original mafic minerals such as clinopyroxene and orthopyroxene are altered to smectite but sometimes they are preserved. Small amounts of carbonate minerals are found in this zone. A zonal sequence of opaque minerals is also found. Pyrite is found in zones (1) and (2). Very tiny amounts of pyrrhotite coexisting with pyrite are found in zone (1). Magnetite is common in zone (3) but this mineral is thought to be original. Fluid inclusion studies have been carried out on quartz samples from the veinlets in the Yugashima Group and diorite porphyry and from the A u - A g vein. Homogenization temperatures of inclusion fluids in the quartz coexisting with laumontite and stilbite range widely from ~240°C to 380°C as shown in Fig. 1.76. This wide range of homogenization temperatures and the coexistence of vapor- and liquid-rich fluid inclusions in the same quartz crystal suggest that boiling took place when these zeolites and quartz were precipitated. The homogenization temperatures of fluid inclusions in the quartz which is in contact with epidote and prehnite range from ~235 to 285°C. All fluid inclusions in the quartz are liquid-dominated type and vapor homogenized into liquid. The range of temperatures for each alteration zone can be estimated from the following chemical reactions and thermochemical data available for these reactions laumontite
: wairakite + H20
(1-23)
yugawaralite = laumontite + quartz
(1-24)
yugawaralite = wairakite + quartz ÷ H20
(1-25)
stilbite
= laumontite + 3 quartz + H20
(1-26)
heutandite
= laumontite + 2 quartz + 3 H20
(1-27)
stilbite
= heulandite + 3 quartz + 3 H20
(1-28)
The coexistence of laumontite and wairakite is common in zone (1). If the saturated water vapor pressure is equal to 0.3 of total pressure (Zeng and Liou, 1982), the temperature for equilibrium reaction (1-23) and saturated water vapor pressure are estimated to be ~170°C and 230 bar, respectively (Liou, 1971b). Zeng and Liou (1982) have shown that yugawaralite is stable at less than ~230°C and a total pressure of 500 bar, under the condition of quartz saturation. However, if the activity of SiO2 is not unity, the boundary for reactions (1-24) and (1-25) may shift to lower temperatures. Liou (1971a) studied the equilibrium for reaction (1-26) and showed that the equilibrium
Chapter I
106 o o
0 ¢- "0 ~
°o.
o ¢:.-
v
a
o
o
-8 &
o r-
~0 = <
U~
r'-
"E
=
LZ
2 ~ ~
-800 • 700
O cO
.~0 i~ o -600
"~
n -500
g
-400 • 300 "200
;i . . . . . . ~ ~ . , [ ~ v' :, O ooo,
=~
0
N
;
. .
o:: ooo"
oo%oi ~G e.= p' e,ooo o o 0 0>3 oooo;. :,,. ~~ eoe ee el, ee, ~,ee= q o ~ o e o poo J e 13 (/} O e ~ e e *4e e e e e . . . . .
• loo
o
AdvancedArgiIlic AIteratlon
o i
200 300 400 ( T ' C ) ~ Zeolile Zone EpidoleZone ~
Au-AgVein ,
,
L'1~" ,
I I
. . . .
.£ -6~'. ............
- 100
-200
"0 o : 5 , . . . . . . . . . . . . . O. 0 Advanced [~] Argillic Alteration ] Montrnorillonite ]
Zone Zeolile Zone
] Epidole Zone
[ ~ Au-Ag Vein Figure 1.76. Schematic c o l u m n at section o f the Seigoshi A u - A g mine area, central J a p a n (Shikazono, 1985a).
temperature and pressure are 170-185°C at 2 - 5 kbar respectively. However, on the basis of experimental studies (Liou, 1971b; Maruyama et al., 1983) it is safe to expect that stilbite was formed under lower temperature conditions, probably less than ~150°C. Thus, it is considered that zone (3) was formed under temperatures less than ~ 150°C. In summarizing the fluid inclusion studies and stability of zeolite minerals, the most likely temperature range o f zone (1), (2) and (3) is estimated to be ~ 2 5 0 - 2 8 0 ° C , 150-230°C and < 150°C, respectively. Boiling of fluids for zones (2) and (3) suggests that the depth for the zeolite zone is probably less than 500 m from the surface and the epidote zone is more than 500 m.
Miocene-Pliocene Hydrothermal Ore Deposits
107
+2 +1 0
A
Broadlands . . t~ \(wel~ll) . _, . - " .~ • Mexixali (well5) Larderello \ \. . ~,; ~ ~, . . . " Nasu Chausudake ,, - , % Wairakei tlveragera- .%k\, ,\ ,,o . ' , ~.-,, • . . . . c ' Showashinzan ~ 0 * {average) :~:4[~ ~"
The G e y s e r s
~-~ 1 O 2
Showashinzan -'~,~, ;*" " '~ . ' " .~'./~huachapan .." - " El Tatio . -
3
Satsuma Iwojirna
...~':i~i~'~ : ::~
\B
-4
50
200
2;0
3;0
3;0
4;0
450
T(°C)
Figure 1.77. Range of hydrogen sulfide fugacity (fH2S) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica-rich zone) and some active geothermal systems. A = propylitic alteration; B = advanced argillic alteration. Data on active geothermal systems are taken from Ellis and Mahon (1977). The calculation of fH2s of Showashinzan for saturated water vapor pressure condition was based on the analytical data on volcanic gas by Matsuo (1961). The calculation of fH2S of Satsumalwo-Jima was based on the analytical data on volcanic gas by Matsuo et al. (1974) assuming PH20 = 0.5 kbar. ThermochemicaI data necessary for estimating fH2s for the propylitic and advanced argillic alterations are taken from Bird and Helgeson (1981) and Giggenbach (1981). The curves A-A', B-B' and C-C' represent the equilibrium of epidote(xp~ =0.30) - - K-mica(aK mica=0.9) - K-feldspar(a~ feidspar=0.95) -- pyrite - calcite - chlorite(aFco_0.5) where Xpis, aK-mica, aK-feldspar, and aFeo are mole fraction of pistacite component in epidote, activity of K-mica component in mica, activity of K-feldspar component in K-feldspar and activity of FeO component in chlorite, and hematite + liquid sulfur .~- pyrite + H2S, respectively (open circle = vapor-dominated system; solid circle = hot-water-dominated system; solid triangle = volcanic gas) (Shikazono, 1985a).
Ranges of $2, 02 and H2S fugacities were estimated as shown in Figs. 1.81, 1.82 and 1.77, respectively. $2 and 02 fugacities were estimated from the alteration mineral assemblage and homogenization temperature data. H2S fugacity was estimated from the chemical compositions of epidote and chlorite following the procedure by Giggenbach (1980, 1981) as shown in Fig. 1.77. CO2 fugacity can be inferred from the following chemical equilibrium relations (Bird and Helgeson, 1981): 3K-mica + 4calcite + 6quartz = 2clinozoisite + 4K-feldspar ÷ 4CO2 + 2H20
(1-29)
Based on the analytical data of K-mica, epidote and K-feldspar and using thermochemical data on these minerals (Helgeson and Kirkham, 1974; Helgeson et al., 1978; Bird and Helgeson, 1981), the fco2 range for the propylitic alteration was estimated (Fig. 1.78). The regional distribution of advanced argillic alteration in this district is shown in Fig. 1.73. The alteration zone of this type is distributed with a given trend, in general running from north to south. The distribution area of this type of alteration is more restricted than that of the propylitic alteration (Fig. 1.73). This type of alteration is well observed in the Funabara area, underground in the Seigoshi mine and the Ugusu silica mine (Figs. 1.73 and 1.79). The original rocks for this alteration are different in these
Chapter 1
108
T.Alfina I */Bagn/re @4
1
0 110
0
0 _1
-1
Broadlands /~" Nilan/d~Krafla (~zildere (~Ngawha ~ ~ 2 p h a s e Rotorua V O~.~J'~
~ ~'~'~'~~'~'r ~
"-~L~------''~2Krafl pha%e ~ \ "CerroPrieto Sa,,onSea
/ ~ ~ A '
,~,and
" Seigoshi
-2 -3
-4 I
[
I
I
150
200
250
300
T(°C)
Figure 1.78. Range of carbon dioxide fugacity (fco2) and temperature for the propylitic alteration (epidote zone) in the Seigoshi area and some active geothermal systems. Seigoshi = propylitic alteration of the Seigoshi district. Data on active geothermal systems are taken from Helgeson (1967, 1968), D'Amore and Panichi (1980), Giggenbach (1980, 1982), Bird and Norton (1981), and Arndrsson and Gunnlaugsson (1983). ThermochernicaI data necessary for estimating fco2 for the propylitic alteration are taken from Helgeson and Kirkham (1974), Helgeson et al. (1978), and Bird and Helgeson (1981). The curves A-B and A ' - B ' are equilibria for epidote(xp~=o._~o) - - K-miCa(aK,.~c~ o.9) - - K-feldspar(aKqeid~p~r-0.95) - - calcite assemblages at saturated water vapor pressure condition.
areas; Tanaba andesitic rocks (Funabara area), basic intrusive rocks, pyroclastic rocks of the Yugashima Group and Tanaba Andesite (Seigoshi area) and Koshimoda Andesite, and andesitic lava and pyroclastic rocks of the Yugashima Group (Ugusu area). The elevation level of this alteration is high, ~200-400 m higher than the top of the A u - A g vein. Forms of the alteration halo are also different in these three areas; lenticular and mushroom like (Funabara), vein (Seigoshi) and strata-bound and lenticular (Ugusu). Among these three areas, the most intense alteration has occurred in the Ugusu area. Thus the description on the alteration in this area is given below as an example of the advanced argillic alteration. The Ugusu silica mine, situated 4 km south of the Seigoshi mine (Fig. 1.73), produces 6 x 105 tons of silica ore annually, which contains more than 95 wt% SiO2. The geology and the lateral and vertical alteration zonings of this mine are shown in Figs. 1.79 and 1.80. Quartz is the most predominant phase in the central zone and the silica content of this zone is 95-99 wt%. Native sulfur is found in the pores of the highly silicified rocks together with small amounts of topaz. Alunite occurs abundantly, surrounding this zone, with quartz. Dickite, sericite and sericite/montmorillonite mixed-layer mineral occur in the more peripheral zone. Montmorillonite (tri-type, contrasting with di-type of the propylitic alteration) occurs in the most peripheral zone. Generally, the boundary between the alunite zone and surrounding clay zone is sharp, but the boundary between the silica zone and alunite zone is gradual. Pyrite occurs in the clay zone. Hematite
Miocene-Pliocene Hydrothermal Ore Deposits /i 400 ~ ',--/-'.~,; - 5OO
.
/
)\
,,/
109
...~'~'--.o~ Hakko
.
-
"x.J.
..
.....6~a.'~::
:ii:"'=ii "'':.
BB
::::: ::..:::
Silica-rich Zone Alunite-rich Zone
L
j
~
1kin
Clay-rich Zone
~2~ Unconformity
Figure 1.79. Geology and alteration zoning in the Ugusu silica mine (modified from Iwao, 1949, 1962). Toi F = Toi Formation; Koshimoda = Koshimoda andesite; Hakko = Hakko orebody; Shibayama = Shibayama orebody. Numbers indicate metres above sea level. Alteration zoning in the section of A'-B ~ is shown in Fig. 1.80 (Shikazono, 1985a)
Margin(A') Lower
Montmorillonite a Cristobalite Ser./Mont
.
Sericite Kaolin Alunite Topaz Native Sulfur Quartz Pyrite Hematite Rutile
.
.
.
.
.
Center(B') Upper
.
........ ___ ___
................ .................
Figure 1.80. Zonal sequence of the advanced argitlic alteration from the central to marginal zone in section of A'-B ~ in Fig. 1.79 and from upper horizon to lower horizon (Shikazono, 1985a)
o c c u r s as d i s s e m i n a t i o n a n d v e i n l e t s w i t h q u a r t z in t h e silicified a n d a l u n i t e zones. R u t i l e c o m m o n l y o c c u r s in the a l t e r a t i o n z o n e , e s p e c i a l l y in the c e n t r a l z o n e . A n a t a s e t e n d s to o c c u r in t h e m a r g i n a l p a r t o f the a l t e r a t i o n z o n e . S m a l l a m o u n t s o f o~-cristobalite a n d t r i d y m i t e o c c u r in the m a r g i n a l z o n e . T h e fluid i n c l u s i o n s c a n b e d i v i d e d i n t o t w o types: v a p o r - a n d l i q u i d - r i c h fluid i n c l u s i o n s . T h e filling d e g r e e o f fluid i n c l u s i o n s f r o m s o m e s a m p l e s f r o m the silicified a n d a l u n i t e z o n e s is v a r i a b l e a n d h o m o g e n i z a t i o n t e m p e r a t u r e s v a r y widely. T h i s i n d i c a t e s
Chapter 1
110
that the liquid-vapor separation occurred during the hydrothermal alteration process for the silica and alunite zones. Although the filling degree is variable, the homogenization temperature of the samples from the silica-rich zone is high, being in the range of 285-430°C; that from the alunite-rich zone in the range of 240-360°C and that from the alunite-clay-rich zone in the range of 220-280°C. This indicates that a steep temperature gradient existed in the hydrothermal alteration zones probably due to the mixing of high temperature volcanic gas and low temperature groundwater. Based on the hydrothermal alteration mineral assemblages and the fluid inclusion, the probable range of gas fugacities (fs2, fOR, fH2S) and temperature can be seen in Figs. 1.81 and 1.82: these estimated fugacities are quite different from those of the propylitic alteration. The characteristic features and differences of the two types of alteration are schematically summarized in Fig. 1.76, indicating the vertical changes in geology, alteration minerals and fluid inclusion characteristics in the Seigoshi-Ugusu district. The vertical alteration zoning from the shallower portion to the deeper portion in this district is summarized as follows: a massive part of the advanced argillic alteration zone; the zeolite zone of the propylitic alteration; and the epidote zone of the propylitic alteration. The advanced argillic alteration zone is developed near the unconformity boundary between relatively permeable pyroclastic rocks and the overlying relatively impermeable pyroclastic rocks. The upper portion of this district (zeolite and advanced argillic zones) is considered to be a two-phase (vapor and liquid) separation zone. It is worth noting that the mineralogical sequence in some active geothermal areas characterized by high chloride concentrations is similar to that found in the propylitic alteration of this district. For example, the alteration zoning in Wairakei, New Zealand, is similar to that of the Seigoshi district: in the shallower zone montmorillonite, illite-montmorillonite mixed-layer mineral, and laumontite occurs, and in the lower part
s~ c'd
co
f
s ~. "" j..-z.----~_ ~3,
-15 -20 -25
150
200
250
300
T (°C)
Figure 1.8i. Range of sulfur fugacity (.Ds2) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica- and alunite-rich zones (Shikazono, 1985a)). A = propytitic alteration; B = advanced argillic alteration. Thermochemical data on the fs2-temperature boundaries for the equilibria of: liquid sulfur ~- S2gas; hematite + pyrite ~- magnetite + S2gas; and pyrite ~ pyrrhotite + S2gas were taken from Heigeson (1969) and Rau et al. (1973a,b). S(1) = liquid sulfur; S2(v) = S 2 g a s ; ht = hematite, m t = magnetite; py = pyrite; po = pyrrhotite.
Miocene-Pliocene Hydrothermal Ore Deposits
tll
•
i.i i-
-30 -35 -40 04 O
"o
-45 -50 -55 -60
I
I
I
I
150
200
250
300
T (°C)
Figure 1.82. Range of oxygen fugacity (f%) and temperature for the propyiitic alteration (epidote zone) and the advanced argillic aIteration (silica- and alunite-rich zones (Shikazono, i985a)). A = propylitic alteration; B = advanced argilIic alteration. Thermochemical data on the fo2-temperature boundaries for the equilibria of: liquid sulfur + hematite ~ pyrite + H20 for saturated water vapor pressure condition; hematite ~ magnetite + O2; and pyrite + magnetite ~ pyrrhotite + 02, are taken from Helgeson (1969) and Rau et al. (1973a,b). Abbreviations used are the same as in Fig. 1.81. epidote and wairakite are found (Steiner, 1968). Both areas (Seigoshi and Wairakei) are characterized by c o m m o n occurrences of zeolites (laumontite and wairakite) and epidote and small amounts o f carbonates. The Larderello region (Italy) is also characterized by the occurrence of wairakite, laumontite and epidote (Cavaretta et al., 1982). Environmental conditions o f gas fugacities (fHzS, fco2) and temperature at the Wairakei and Larderello regions are similar to those of the Seigoshi district as shown in Figs. 1.77 and 1.78. The estimated depth o f the part affected by the zeolite and epidote alterations in the Seigoshi district is consistent with those of Wairakei and Larderello. The formation of epidote, K-feldspar, prehnite, wairakite and calcite in the geothermal area is considered to be due to the loss o f CO2 gas and rapid precipitation from the solution supersaturated with respect to quartz (Browne, 1978). The widespread occurrence o f these minerals in the Seigoshi district seems to be consistent with the above-mentioned consideration, namely that these minerals usually occur as veinlets rather than the replacements of original minerals and filling amygdule. In particular, many veinlets of epidote, prehnite and wairakite are found near the A u - A g - q u a r t z veins. There are three possible mechanisms for generating the strongly acid solution which caused the advanced argillic alteration: (1) alteration caused by the vapordominated system as inferred by White et al. (1971); (2) alteration caused by the oxidation o f HzS near the surface; and (3) alteration by volcanic gas a n d / o r hot water condensed from a volcanic gas. A m o n g them, (3) is the most attractive mechanism given the following evidence and considerations.
112
Chapter 1
(1) The temperature of the advanced argillic alteration estimated from the fluid inclusion studies and mineral assemblages, varies widely from ~220°C to ~420°C. The homogenization temperature for the central part of the alteration zone is from ~285°C to ~430°C. This temperature range is higher than the temperature of the vapor-dominated system defined by White et al. (1971) who showed that vapor-dominated systems such as Larderello (Italy), The Geysers and Mud Volcano (Yellowstone Park, Wyoming, USA) and Matsukawa (northeast Japan) are characterized by the temperature of around 240°C. (2) The advanced argillic alteration minerals are similar to those of the vapordominated system. But, in general, the minerals indicating relatively higher temperatures, such as pyrophyllite and diaspore, are lacking in the vapor-dominated system mentioned above. However, in Matsukawa which may be a vapor-dominated system at present, these high-temperature minerals like pyrophyllite, diaspore and andalusite have been reported (Sumi, 1968a). However, these minerals are considered to have been formed during the past and are not presently forming (Sumi, 1968b). (3) The minerals containing F and C1 such as topaz and zunyite are common in the advanced argillic alteration. The HF activity of the solution was estimated to be ~0.01 based on the F content of topaz from the Ugusu mine (Shibue and Iiyama, 1984). This high F concentration is observed in Crater Lake (Mt. Ruapehu, New Zealand) (Giggenbach, 1974) and is interpreted to be due to the injection of acid fumarolic gases to the lake (Giggenbach, 1974). Gases collected from a fumarole at Showashinzan and Satsumaiwo-Jima also have a high H F / H 2 0 volume ratio which is in the range of (2-5)× 10 -4 (Matsuo et al., 1974; Mizutani and Sugiura, 1982). (4) If alunite, K-mica and kaolinite (which are common minerals in the advanced argillic alteration) are in equilibrium, the concentration of H2SO4 can be estimated based on the experimental work by Hemley et al. (1969); the concentration of H2SO4 at 200°C and 300°C is 0.002 and 0.012 M, respectively. This may suggest that it is difficult to form such a high concentration of sulfate ion only by oxidation of H2S. (5) The fugacity of H2S (fH2s) for the advanced argillic alteration is plotted in Fig. 1.77 together with those for some active geothermal systems and volcanic gas. The fH2S for the advanced argillic alteration was estimated on the basis of hematite-pyriteliquid sulfur equilibrium. Estimated fH2S for the advanced argillic alteration is lower than that for the vapor-dominated system (Larderello and The Geysers) but is similar to that of the volcanic gas collected from one of the fumaroles of the Showashinzan volcano in Japan. Only a few areas subjected to solfataric alteration have been well studied. Satsumaiwo-Jima is probably the only area where alteration caused by volcanic gas has been studied in detail. The zonal sequence from the centre to the margin and from a topographically higher level to a lower level is from a silica-rich zone through an alunite-rich zone to a clay (montmorillonite)-rich zone (Yoshida et al., 1976). This pattern is similar to that of the advanced argillic alteration in the Seigoshi-Ugusu district, although oe-cristobalite is abundant and quartz is poor in the Satsumaiwo-Jima. But in the Seigoshi oe-cristobalite is poor and quartz is abundant. In the Satsumaiwo-Jima pyrophyllite, diaspore, zunyite and topaz are not found, although these minerals occur in the Seigoshi-Ugusu district. The occurrence of alunite in the Satsumaiwo-Jima indicates that the formation of alunite is largely controlled by the existence of groundwater. The
Miocene-Pliocene Hydrothermal Ore Deposits
l 13
alunite zone in this island is strata-bound in form. In the Ugusu mine the occurrence of alunite is very similar. It is interesting to note that the advanced argillic alteration tends to occur near the unconformity boundary (Fig. 1.79). The similarity in the mode of occurrence of alunite from the Satsumaiwo-Jima and from the Ugusu mine, the existence of the unconformity and the steep temperature gradient from central to marginal part in the Ugusu mine may suggest that the mixing of groundwater with volcanic gas and/or hot water condensed from the volcanic gas is the most important mechanism for the formation of advanced argillic alteration also in the Satsumaiwo-Jima area. The coexistence of these two types of hydrothermal alteration and associated high sulfidation-type A u - A g deposits (acid alteration) and low sulfidation-type Au Ag deposits (neutral alteration) is observed not only in the Seigoshi-Ugusu district, but also in the other epithermal A u - A g vein-type mine districts in Neogene volcanic region in Japan: Izu Peninsula, Takadama, Kushikino, Hishikari, Akeshi-Kasuga, Yatani, and Harukiyama A u - A g mine (Hokkaido) (Yamada, 1995) districts. In these districts both the epithermal vein-type deposits and highly silicified and clay-rich rocks are found in close proximity in time and space. The characteristic differences between the two types of coexisting associated alterations in each area are very similar to those in the Seigoshi-Ugusu district. For instance, advanced argillic alteration zone lies at topographically higher levels and in younger formations than the epithermal vein-type deposits and associated propylitic alteration. However, some differences exist in a different area. For instance, in the Kasuga-Akeshi area, fine Au particles occur in the highly silicified rocks. Surrounding the highly silicified rocks, the following zonal sequence is observed from the centre to the margin (Tokunaga, 1955; Saito and Sato, 1978): an alunite zone, kaolinite zone and a montmorillonite zone. This sequence is similar to that observed in the Ugusu deposits, but Au has not been detected in the silicified part of the Ugusu silica deposits. Topaz and zunyite have not yet been reported from the Kasuga silicified rock, while they are found in that of the Ugusu. Homogenization temperatures for the silicifled rocks in the Kasuga, Akeshi, Iwato and Kushikino areas are generally lower than those for the Ugusu deposits (Takenouchi, 1981). Boiling phenomena have been observed in the fluid inclusions from the silicified rocks in these areas (Takenouchi, 1981). Yamada (1995) described acid and potassium alterations and high sulfidation and low sulfidation mineralizations in the Harukiyama district (Hokkaido). He indicated based on K - A r data on the alterations that the high sulfidation-type have been formed first by the residual magmatic fluid subsequent to the intrusion of quartz porphyry magma and then the low sulfidation-type could have been formed by the circulation of meteoric water generated by the heated emanation from hot magma solidified but still similar temporal relationship is observed in the other districts (Osorezan: Aoki, personal communication, 1990; Yatani: Shikazono, unpublished; Seigoshi-Ugusu: Shikazono, unpublished).
1.4.2.6. Chemical composition of alteration minerals Although a wide range of alteration minerals has been recognized in epithermal systems considered here, few of their chemical compositions have been determined. Trioctahedral chlorite occurs commonly in geothermal and hydrothermal areas, whereas the occurrence of dioctahedral chlorite is very limited. For instance, donbasite
114
Chapter I
does not occur in geothermal and hydrothermal areas, but sudoite (Al-chlorite) commonly does not occur in Kuroko mine area (e.g., Tsuzuki and Honda, 1977). Dioctahedral chlorite has not been reported from the Neogene C u - P b - Z n vein-type deposits in Japan; instead, trioctahedral chlorite is common (Shirozu et al., 1975). In Kuroko mine area, the common chlorite is trioctahedral Mg chlorite. The structural formula for trioctahedral chlorite is represented by (Mg6-x-yF@ + Alx)(Alx Si4-x)O10(OH)8. It is convenient to plot a diagram of Fe2+/(Fe 2+ + Mg) (in atomic fraction) against ~VA1/(Si + IVA1)to display compositional variations in chlorite (Fig. 1.83) (e.g., Hey, 1954; Foster, 1962, Nagasawa et al., 1976). Although a large body of analytical data on hydrothermal chlorite from geothermal and hydrothermal areas is available, there are few data on the chemical composition of original fresh volcanic rocks. The relationship between the ratio MgO/FeO in chlorite and that in the fresh host-rocks indicates that the value of MgO/FeO of chlorite generally satisfies a line of 1 : 1 slope. The correlation between MgO/FeO in the host rock and that in the chlorite implies that the MgO/FeO value of chlorite from propytitically altered rocks associated with the mine areas and from altered rocks in terrestrial and submarine geothermal areas is largely affected by MgO/FeO ratio of original fresh rocks. However, most of the data from Kuroko and Neogene C u - P b - Z n vein-type deposits deviate significantly from this line. Chlorite compositions from areas (Toyoha Pb-Zn vein, Kuroko deposits) deviate significantly from a line of 1 : 1 slope. This deviation implies that the Fe2+/Mg value of chlorite from these areas is controlled not only by the FeO/MgO value of the fresh host rocks, but also by factors such as the ratio of Fe 2+ to Mg 2+ in the fluid phase. As noted already, several investigators have shown isotopically that seawater played an important role in the formation of Kuroko deposits (e.g., Sakai et al., 1970; Kajiwara, 1971; Hattori and Sakai, 1979; Farrell and Holland, 1983). Mg-rich chlorite occurs in gypsum-anhydrite bodies in many Kuroko deposits. Farrell and Holland (1983), Shikazono et al. (1983), and Kusakabe and Chiba (1983) suggested the involvement of large amounts of seawater or seawater-dominated hydrothermal solution in the formation of the gypsum-anhydrite bodies. It is reasonable to assume from the large number of experimental studies on rock-seawater interaction at elevated temperatures that the seawater-dominated fluid phase, which interacted with volcanic rocks at a high water/rock ratio, contained appreciable amounts of Mg but very small amounts of Fe (Seyfried and Mottl, 1982). Therefore, it is likely that Mg-rich chlorite precipitated from a solution with a high proportion of Mg 2+ to Fe 2+. In contrast to the Kuroko hydrothermal system, there is no evidence for the involvement of seawater in the hydrothermal system associated with the C u - P b - Z n vein mineralization. Results of hydrogen and oxygen isotopic studies indicate that large amounts of meteoric water were incorporated into the ore fluids responsible for these vein-type deposits (Hattori and Sakai, 1979). If F e - M g chlorite is assumed to be in equilibrium with a fluid phase and pyrite, the ratio of Fe 2+ to Mg 2+ in the fluids may be related to factors such as pH, f Q , temperature and total dissolved sulfur concentration (ES). This relationship can be derived from the following chemical reactions: MgsA1Si3A1Olo(OH)8 + 5 Fe 2+ = FesA1Si3A1Olo(OH)8 + 5 Mg 2+
(1-30)
Miocene-Pliocene Hydrothermal Ore Deposits
115
1.0
oI 03 01
04
0.8
05
._.0.6
o)
o 6
oO6
4LL
o6
v
o6
u_ 0.4
°6o6
o6
0.2
9 ~ 910_10 • •8 o 8 8 o8
09
0.0
018
1:0
1:2 1:4 AI in 4(AI,Si)
o7o 7
1:6
1:8
2~0
Figure 1.83. Variation of Fe2+/(Fe 2+ -}- Mg) and tetrahedral A1 of chIorite from hydrothermal ore deposits: Japanese Neogene Cu-Pb-Zn vein-type (open circle) and Kuroko deposits (solid circle). Localities: 1: Ashio (Nakamura, 1960, 1963); 2: Yatani (Hattori, 1974); 3: Toyoha (Shikazono 1974a, Sawai, 1984); 4: Kishu (Shirozu, 1958); 5: Sayama (Shirozu, 1958); 6: Mikawa (Nagasawa, 1961); 7: Furutobe (Shirozu et al., 1975); 8: Hanaoka (Hayashi 1961, Hayashi and Oinuma, 1965; Tsuzuki and Honda, 1977; Shirozu et aI., 1975); 9: Wanibuchi (Sakamoto and Sudo 1956, Iwao and Minato 1959, Katsumoto and Shirozu, 1973); 10: western Bergslagen (Baker et al., 1983) (Shikazono and Kawahata, 1987).
FeS2 + 2 H + + H 2 0 = F e 2+ + 2H2S + 1 / 2 0 2
(1-31)
w h e r e MgsA1Si3A1OIo(OH)8 and Fe5A1Si3A1Olo(OH)8 represent M g - c h l o r i t e and Fechlorite, respectively. Generally, activities o f liquid H 2 0 and FeS2 do not deviate f r o m unity; thus these values are a s s u m e d to be unity.
116
Chapter 1 From equations (1-30) and (1-31), we obtain, for the H2S dominant region,
log(aFe-chl/aMg-chl) = log K1-30 + 5 log Kl-31 -- 5 log aMg2+ -- 10 log ES - 10 log gHzS -- (5/2)1og fo2 -- 10pH
(1-32)
where g is the activity coefficient. For the SO 2- dominant region,
1og(aFe-chl/aMg-chl) = log K ~_30+ 5 log K 1-31 - 5 log aMg2+ 10 pH -- 10 log ES - 10 log YSO2- + 35/2 log fo2 + 10 log KI-34
(1-33)
where KI_34 is the equilibrium constant for the following reaction, H2S+202=SO ] +2H +
(1-34)
Equations (1-32) and (1-33) imply that the aFe--chl/aMg-chl of chlorite depends on foz, aMgz+, ~S, pH, temperature and ionic strength. From equations (1-32) and (1-33), it can be shown that the aFe-chl[aMg-chlof chlorite in equilibrium with pyrite decreases with increasing fo2 at constant temperature, pH, aMg2+, ionic strength, and ES in the region where H2S is dominant, whereas it increases with increasing f Q in the SO ] - dominant region. Previous studies on the estimates of f Q , fs2, pH, ZS and temperature for Kuroko ore deposition have been reported, for instance, by Kajiwara (1971), Shikazono (1976) and Ohmoto et al. (1983), who showed that the fo2 of the Kuroko ore fluids lies in the region close to the SO42-/H2S boundary (in this expression, SO ] and H2S represent the concentration of the total dissolved oxidized sulfur species and total dissolved reduced sulfur species, respectively). These estimates seem consistent with the tow Fe2+/Mg value of chlorite from Kuroko deposits. Chlorite from the Toyoha Pb-Zn vein type deposits is associated with sphalerite, pyrite and, rarely, pyrrhotite (Shikazono, 1974a). The iron content of the sphalerite associated with chlorite is 1.2-2.9 wt%. The temperature of formation of chlorite in the Toyoha deposits is estimated to be 200-250°C from fluid inclusion data (Shikazono, 1974a, 1975; Yajima and Ohta, 1979), which indicates that the chlorite was formed in a relatively reducing environment in which reduced sulfur species predominate. This estimate seems consistent with the composition of chlorite from the Toyoha deposits and from the other Neogene C u - P b - Z n vein-type deposits in Japan, which contains up to 40 wt% FeO (Nakamura, 1960, 1963; Shirozu, 1969; Shikazono, 1974a; Hattori, 1974). Shikazono (1974a,c, 1977a, 1978b) and Hattori (1975) showed that these deposits formed in the environments where species of aqueous reduced sulfur predominated. This estimated range of oxidation state and ratio of concentration of aqueous reduced sulfur species to oxidized sulfur species appears to be in agreement with the chemistry of chlorite from these deposits. However, the Fe2+/Mg of chlorite also depends on the other factors, such as IES, pH and aMg2+. From equations (1-32) and (1-33), it is obvious that increasing IgS and aMg2+ also causes a lower Fe2+/Mg in chlorite. The Fe 2+ /Mg and Fe-3+ /Fe 2+ values of chlorite from Kuroko deposits and Neogene C u - P b - Z n vein-type deposits differ greatly (Fig. 1.83). Chlorite from Kuroko deposits contains lower Fe2+/Mg and higher Fe3+/Fe 2+ values than this from the Neogene vein-type deposits in Japan. The most likely explanation for these differences is that these two types of deposit formed at different states of oxidation, although other
Miocene-Pliocene Hydrothermal Ore Deposits
117
variables such as pH, temperature, aMg2+ and NS are also possible important factors, as discussed above. Generally, the Fe z + / M g value of chlorite from a given geothermal or hydrothermal area is variable. The Fe2+/Mg value of chlorite in the Toyoha mine district varies widely. Sawai (1984) has shown that the iron content of chlorite away from the P b - Z n veins. Iron content of chlorite in the Toyoha veins is very high (40 wt% FeO: Shikazono, 1974a; Sawai, 1984). Variations in iron content of chlorite from the host rocks toward the C u Pb-Zn veins have also been studied for other Neogene vein-type deposits in Japan (e.g., Ohe, Ashio deposits: Hayashi, 1979). Variations in iron content of chlorite from Japanese epithermal mine districts suggest that the iron content of chlorite in the discharge zone of a hydrothermal system is higher than that in the recharge zone. Mottl (1983) has also suggested that the FeZ+/Mg value of chlorite in the discharge zone of submarine geothermal systems is higher than that in the recharge zone. The Fe2+/Mg of chlorite from different parts of a geothermal system is higher than that in the recharge zone. The FeZ+/Mg of chlorite from different parts of a geothermal system (recharge zones versus discharge zones) is probably not constant. In order to evaluate the effect of the fluid movement at discharge versus recharge zones, the temperature dependence of Fe2+/Mg 2+ in fluids in equilibrium with the isochemically recrystallized crystal rocks will be considered below. By taking a value for Fe2+/Mg of chlorite that is equal to that of the average andesitic and basaltic rocks (0.6), and assuming that chlorite is an ideal solid solution of 14 A Fe-chlorite and 14 A Mg-chlorite, the dependence of a aFez+/aMg2+in fluids on temperature was calculated by using thermochemical data for chlorite from Walshe and Solomon (1981). Figure 1.84 shows that aFez+/aMg2+of fluids increases with increasing
-2
+-3 04
~-4 o
200
250
3()0 350 Temp.(°C)
Figure 1.84. Variation of aFe2+/aMg2+of hydrothermal solution in equilibrium with chlorite having constant Fe2+/Mg (=0.6), as a function of temperature. The significance of points A and B is discussed in the text (Shikazono and Kawahata, 1987).
118
Chapter I
of temperature. Therefore, if fluids initially in equilibrium with chlorite having the same FeO/MgO as that of average andesitic and basaltic rock at elevated temperature (for example, at point A in Fig. 1.84) ascend rapidly without interaction with the surrounding rocks, chlorite precipitating from fluids at lower temperature (for example, at point B in Fig. 1.84) could contain appreciable amounts of Fe 2+ compared with Mg. This mechanism could lead to the formation of chlorite having an unusually high content of iron. It is also noteworthy that the vein chlorite in the altered basalt from the Costa Rica Rift contains higher concentrations of iron than the chlorite that replaces mafic minerals in the rock (Kawahata, 1984). In this case it is likely that the flow rate of ascending fluids from which vein chlorite precipitated is high compared with the rate of reaction between the ascending fluids and surrounding rocks. The Fe2+/Mg value of chlorite precipitating from ascending fluids depends on the extent of deviation from equilibrium between fluids and surrounding rocks. As discussed in detail by Giggenbach (1984), a number of processes such as adiabatic and conductive cooling of fluids and mixing of fluids can cause this deviation. The above considerations suggest that chlorite occurring in the discharge zones of hydrothermal systems would contain a higher concentration of iron than that occurring in recharge zones. If a magnesium-rich solution such as seawater or a seawater-dominated hydrothermal solution was involved in the hydrothermal systems, it could be expected that chlorite, even that occurring in discharge zones, would contain a high content of Mg. It is widely accepted that ascending hydrothermal solutions mixes with cold seawater at the time of ore formation at Kuroko (e.g., Hattori and Sakai, 1979; Shikazono et al., 1983). The Mg concentration of this hydrothermal solution increased with higher degrees of mixing. Involvement of large amounts of seawater at the site of ore deposition could be one of the reasons why chlorite from Kuroko deposits contains high amounts of Mg. In addition to a large concentration of Mg in these fluids, relatively high fo2, NS and SO42-/H2S, and low pH, could lead to the large Mg content of chlorite in Kuroko deposits, as discussed above. Consequently, the composition of chlorite in the discharge zone depends largely on the chemical nature of fluids (factors such as Fe2+/Mg 2+, SO42-/H2S, pH, aMg2+) and temperature. Movement of fluids may also be an important cause for the variability in the ratio of Fe 2+ to Mg in hydrothermal chlorite. Wide compositional variations in chlorite from the hydrothermal ore deposits in Japan, including Kuroko and Neogene C u - P b - Z n vein-type deposits, are considered to reflect the variable chemical nature of ascending ore fluids and fluids that mix with ascending ore fluids at discharge zone. The variations in Fe and Mg contents of the 14 ,~ Fe-chlorite-14 ,& Mg-chlorite solid solution are considered here. However, structural formulae for chlorite are not as simple as those considered here. As mentioned by Walshe and Solomon (1981), Stoesell (1984), Cathelineau and Nieva (1985) and Walshe (1986), chlorite solid solution may be represented by six components, and accurate thermochemical data on each end-member component at the hydrothermal conditions of concern are necessary to provide a far more rigorous calculation of the equilibrium between chlorite and hydrothermal solution. However, the above argument demonstrates that the composition of chlorite is a highly useful indicator of physicochemical conditions of hydrothermal solution and extent of water-rock interaction.
Miocene-Pliocene Hydrothermal Ore Deposits
119
Epidote is one of the most common alteration minerals occurring in geothermal and mine areas. The factors controlling the chemical composition of epidote from geothermal areas have been examined from thermochemical points of view. Bird and Helgeson (1981) discussed the effects of temperature, CO2 and 02 fugacities, and activity ratios (aFe3+)/(aH+) 3, (aca2+)/(au+) 2, and (aai3+)/(aH+) 3 in geothermal waters on the variations in iron and aluminum contents of epidote. Giggenbach (1981) calculated the effects of partial pressure of CO2 gas and temperature on iron content of epidote in equilibrium with alteration minerals such as kaolinite, K-mica, K-feldspar, albite, paragonite and calcite in geothermal systems. D'Amore and Gianelli (1983) showed the dependence of iron content of epidote on 02 fugacity for the epidote-K-feldspar-albitetremolite-chlorite equilibrium assemblage. Wolery (1978) and Reed (1982, 1983) have indicated based on a computer calculation of the change in chemistry of aqueous solution and mineralogy during seawater-rock interactions that epidote is formed under the low water/rock ratio less than ca. 50 by mass. Humphris and Thompson (1978), Stakes and O'Nell (1982) and Mottl (1983) have also suggested on the basis of their chemical and oxygen isotopic data of the altered ridge basalts that epidote is formed by seawater-basalt interaction at elevated temperatures (ca. 200-350°C) under the rock-dominated conditions. If epidote can be formed preferentially under such low water/rock ratio, the composition of epidote should be influenced by compositions of the original fresh rocks. Shikazono (1984) summarized analytical data of the epidote from geothermal areas to consider the relationship between the composition of epidote and that of the original fresh rocks and to inspect the other factors controlling the compositional variations in epidote. The discussion on the epidote composition by Shikazono (1984) is described below. Chemical compositions of epidote and original rocks in several geothermal and mine areas including Seigoshi (Japan, epithermal Au-Ag vein-type mine area), Yugashima (Japan, ancient geothermal area), Furutobe (Japan, Kuroko mine area), Ohtake (Japan, active geothermal area), Mid-Atlantic ridge, Costa Rica rift, Mitsuishi (Japan, pyrophyllite deposits), Shimokawa (Japan, ancient ophiolite area associated with Besshitype deposits), Reydarfjordur (Iceland, ancient geothermal area), Kushikino (Japan, epithermal Au-Ag vein-type mine area), Hachimantai (Japan, active geothermal area), Broadlands (New Zealand, active geothermal area), Wairakei (New Zealand, active geothermal area), Larderello (Italy, active geothermal area), and Sarmiento ophiolite complex (Chile, ancient ophiolite suite). In individual districts, these rocks are generally subjected to intense hydrothermal alteration. The original rock composition was estimated from analytical data of fresh rocks in the same area. Figure 1.85 shows that the Fe203 content of epidote positively correlates to Fe203 content of original rocks, although in general Fe203 content of epidote from each geothermal area has a variation over several wt%. Therefore, it is inferred that Fe203 content of original rocks has a large influence on iron content of epidote. Original rocks in Fig. 1.85 are mainly andesitic and basaltic rocks of island arcs and ocean ridge. In Fig. 1.85 iron contents of epidote from two different geologic environments, island arc and oceanic ridge or ophiolite, are summarized. It can be seen in Fig. 1.85 that the iron content of epidote from ridge basalt and ophiolite is generally lower than
Chapter 1
120
S
M. q
~13 O
2'
u..
1-
d"
sht
F,L=--~ Y
M, o ,cT J
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 Fe20s Content of Epidote (weight
%)
Figure 1.85. Relation between Fe203 content of epidote and that of original fresh rocks. S: Seigoshi, Y: Yugashima, N: Noya, F: Furotobe, O: Ohtake, M: Mid-Atlantic ridge, C: Costa Rica rift, Mi: Mitsuishi, Sh: Shimokawa (Shikazono, 1984)
that from the volcanic rocks of island arcs. One of the most likely explanations for this difference is the difference in Fe203 content of original rocks; Fe203 contents of oceanic basalts and ophiolite are generally lower than those of island arc volcanic rocks. The positive correlation shown in Fig. 1.85 is consistent with the considerations on seawater-volcanic rock interactions under the hydrothermal conditions by Wolery (1978), Humphris and Thompson (1978), Reed (1982, 1983), Stakes and O'Neil (1982) and Mottl (1983). They suggested that epidote is formed under a low water/rock ratio less than ca. 50 by mass. Mineralogical changes resulted from seawater-volcanic rocks interactions have been elucidated by these recent theoretical studies. The epidotes except those from the Furutobe Kuroko mine area, Mid-Atlantic ridge, Costa Rica rift, and Shimokawa are suspected to have been formed in subaerial geothermal environments. Thus it is uncertain in these cases that the epidote tends to form at a low water/rock ratio. However, it seems likely that the Fe/A1 ratio of epidote formed in subaerial geothermal areas is also reflected by that of original rocks because the concentrations of Fe 3+ and A13+ in geothermal waters interacted with the rocks under the hydrothermal condition are generally very low except for strongly acid and high chloride solutions. Iron content of epidote depends on kind of iron minerals coexisting with epidote. The iron content of epidote coexisting with hematite is relatively high. For example, the epidote coexisting with hematite from the Yugashima, Furutobe and Reydarfjordur has higher iron content than that from the other districts. This relation holds even in geothermal systems (Salton Sea: Keith et al., 1968; Furutobe: Shikazono et al., 1995). On the other hand, the epidote coexisting with pyrite and pyrrhotite contains generally smaller amounts of iron as observed in the Larderello (Cavaretta et al., 1982), Shimokawa, MidAtlantic ridge, and Costa Rica rift. This difference in the iron content of epidote coexisting with different opaque minerals could be explained by the difference in oxygen fugacity. The iron content of epidote coexisting with prehnite from the Seigoshi district is smaller than that without prehnite. It has been clarified that prehnite is stable at the low
Miocene-Pliocene Hydrothermal Ore Deposits
121
CO2 fugacity and that the iron content of epidote in equilibrium with prehnite is lower than that in equilibrium with other minerals such as K-feldspar, K-mica and calcite under such low CO2 fugacity conditions (e.g., Cavaretta et al., 1982). Therefore, it seems clear that iron content of epidote is affected also by CO2 fugacity. Although Fe203 content of epidote from a given geothermal area has roughly a positive correlation to that of original rocks, the variation in Fe203 content of epidote from a given geothermal area ranges over several weight percents. This variation is too large to be explained only in terms of the difference in Fe203 content of original rocks because the composition of epidote varies sometimes so widely even in a single grain. Compositional zoning in epidote grains from subaerial and island arc geothermal systems is commonly known; usually, iron content of epidote increases from core to margin in a crystal. This tendency is observed in the Seigoshi, Furutobe, Yugashima, and Larderello areas, even though a decrease in the iron content from core to margin is rarely recognized in a single grain from the Larderello (Cavaretta et al., 1982). Complicated compositional zonings of epidote from the Salton Sea geothermal area are also found (McDowell and McCurry, 1978). In contrast, epidote from midoceanic ridge basalts (Melson and Van Angel, 1966; Humphris and Thompson, 1978; Kawahata, 1984) shows no wide compositional variation in a single grain. Epidote from a typical midoceanic ridge basalt has a small variation range of iron (Humphris and Thompson, 1978). The variation range in epidote from the Costa Rica Rift studied by Kawahata (1984) is wider than that observed by Humphris and Thompson (1978). Kawahata (1984) considered that the alteration minerals including epidote in the DSDP Hole 504B in the Costa Rica rift were formed by the interaction of midoceanic ridge basalt and ascending hot fluids. The basalt in the Shimokawa area is probably an abyssal tholeiite (Kohsaka, 1975). In this area, strata-bound cupriferous pyritic deposits are found and general features of these deposits resemble those of the ore deposits of Cyprus type and the ore deposits found in East Pacific Rise 21°N. Therefore, the wide variation range in epidote from this area may be explained by the effects of the chemical compositions and gaseous fugacities of ascending hot fluids responsible for the ore formation together with the chemical compositions of original basalt. The compositional zoning of an epidote grain and wide compositional variations in epidote grain from each geothermal area are difficult to explain because iron content of epidote depends on many factors other than the chemical composition of original rocks. However, the differences in the variation range and compositional zoning of epidote from the subaerial and oceanic ridge geothermal systems suggest that the change of physicochemical variables such as oxygen fugacity and temperature during the hydrothermal activities in both geothermal systems is different. Few data on the chemical compositions of feldspars (albite, K-feldspar) are available. Fujii (1976) indicated that K-feldspar and albite in the propylite of west Izu Peninsula, middle Honshu are of nearly end member composition. Nagayama (1992) showed that K-feldspars in the Hishikari A u - A g vein and in the host andesitic rock have different composition; N a / K ratio of K-feldspars from the vein is lower than that from the host rocks.
Chapter 1
122
1.4.2.7. Causes for hydrothermal alteration Hydrothermal alteration is reflected by the changes in many variables (temperature, water/rock ratio, extent of water-rock interaction (reaction progress), reaction rate, flow rate of fluids etc.) (Fujimoto, 1987). Theoretical and experimental works on hydrothermal alteration were reviewed by Meyer and Hemley (1967), and Rose and Burt (1979). In the last two decades, great progress has been made in the field of hydrothemal alteration studies, mainly from computation works on water-rock interactions at elevated temperatures (e.g., Wolery, 1978; Reed, 1983, 1997; Takeno, 1989). These studies revealed the relationship between the changes in chemical composition of hydrothermal solution and the relative abundance of minerals in the rocks. There are different approaches to the study of hydrothermal alteration. For instance, Shikazono (1978a) showed the relationship between chemical composition of hydrothermal solution in equilibrium with the alteration minerals and C1- concentration in hydrothermal solution. Giggenbach (1984) calculated the effect of temperature on the chemical composition of fluids buffered by alteration minerals. The causes for the hydrothermal alteration considered below are mainly based on the works by Shikazono (1978a) and Giggenbach (1984). The effect of the extent of water-rock interaction is not taken into account. Figure 1.86 illustrates the variations in the chemical composition of chloride-rich hydrothermal solution in equilibrium with common alteration minerals with temperature. Figure 1.86 demonstrates that (1) the chemical compositions of hydrothermal solution
B
Na
H
E~ 0
Ca 2.
----4. i
-1
A
~
,-
I
-2
-3
~
150
200
r
250
t
300
Temp.(°C ) Figure 1.86. Variation in chemical compositions (in molaI unit) of hydrothermat solution with temperature. Thermochemical data used for the calculations are from Helgeson (1969). Calculation method is given in Shikazono (1978a). Chloride concentration in hydrothermal solution is assumed to be i moI/kg H20. A-B: Na + concentration in solution in equilibrium with low albite and adularia, C-D: K + concentration in solution in equilibrium with low albite and adularia, E-F: H4SiO4 concentration in equilibrium with quartz, G-H: Ca 2+ concentration in equilibrium with albite and anorthite (Shikazono, 1978a, 1988b).
Miocene-Pliocene Hydrothermal Ore Deposits
123
depend on alteration minerals, temperature and C1- concentration; (2) K + and H4SiO4 concentrations increase with an increase in temperature, while Na + concentration does not largely depend on temperature. Calculations were made, assuming that the dominant anion is C1-. From these calculations, it is considered that the potassic alteration occurs when hydrothermal solution initially in equilibrium with propylitic alteration minerals ascends rapidly and interacts with country rocks at lower temperature ( I - J - K in Fig. 1.86). In this case, addition of K + to the rock takes place. K-bearing minerals such as sericite and K-feldspar precipitate from the fluid accompanied by the destruction of plagioclase in the country rocks and liberation of Ca, Sr, and Na to the fluids. Dissolutions of anorthite and albite components in a plagioclase occur by this mechanism. Thus, it is likely that Ca, Sr and Na are extracted from the rocks. It is expected that SiO2 content of the country rocks increases with progressive alteration because solubility of SiO2 decreases with a decrease in temperature ( O - P - Q in Fig. 1.86) (e.g., Holland and Malinin, 1979). If fluids initially in equilibrium with quartz ascend rapidly, some metastable minerals (amorphous silica, cristobalite, wairakite) may precipitate because of supersaturation with respect to SiO2 (e.g., Wolery, 1978; Bird and Norton, 1981). Important processes for the supersaturation and deviation from the equilibrium between fluids and rocks are adiabatic boiling, mixing of fluids and conductive cooling of fluids (Giggenbach, 1984). Formation of albite which is characteristic mineral of propylitic alteration occurs by heating of rocks and descending fluids at recharge zone in the hydrothermal system (Giggenbach, 1984; Takeno, 1989). Thus, it is considered that the propylitic alteration takes place at recharge zone in the hydrothermal system, while potassic alteration at discharge zone. Besides the effect of temperature, boiling and degassing play an important role for potassic alteration. Degassing of CO2 is accompanied by an increase in pH, causing the depositions of adularia and calcite (Browne, 1978). Probably, formations of some zeolite minerals (wairakite, yugawaralite, taumontite) are also caused by degassing and boiling of fluids and increase in pH (Shikazono, 1985a). There are two important chemical reactions which cause intermediate argillic and advanced argillic alterations H2S + 202 = 2H + + SO]
(1-35)
4SO2 + 4 H 2 0 = H2S + 3 H2SO4
(1-36)
Oxidation of H2S (reaction (1-35)) occurs under the near-surface environment. Oxygen may be supplied from oxygenated groundwater. These oxidation reactions liberate H + ion, leading to a decrease in pH. Under low pH conditions intermediate argillic alteration minerals (e.g., kaolinite, sericite) are stable. When temperatures of volcanic gases containing SO2 decrease, the reaction (1-35) proceeds to the right hand side. This reaction causes a considerable decrease in pH due to the formation of sulfuric acid. Advanced argillic alteration is formed by the interaction of volcanic gas with groundwater. The above interpretation on the alteration zoning is mainly based on thermodynamics. However, it is necessary to consider the influence of kinetics and fluids flow on
124
Chapter 1
the hydrothermal alteration processes to interpret the precipitations of metastable phases such as cristobalite. The coupled precipitation kinetics-fluid flow model was applied to the distribution of SiO2 content and K20 content of the hydrothermally altered andesite in the Hishikari A u - A g mine area, south Kyushu, Japan by Shikazono et al. (2002). This will be described in section 1.4.6. 1.4.3. Geochemical characteristics
Numerous geochemical data (fluid inclusions, stable isotopes, minor elements) on the epithermal vein-type deposits in Japan are available and these data can be used to constrain geochemical environment of ore deposition (gas fugacity, temperature, chemical compositions of ore fluids, etc.) and origin of ore deposits.
1.4.3.1. Fluid inclusions Substantial amounts of homogenization temperature data on the Neogene vein-type deposits in Japan are available (e.g., Enjoji and Takenouchi, 1976; Shikazono, 1985b) and they are summarized in Fig. 1.87. Homogenization temperatures vary widely within a given deposit type and even within a single deposit. However, the range of homogenization temperatures differs according to the type of deposit: 190°C 145 250oc for Au-Ag-rich deposits, 200°C to 250°C for Pb-Zn-Mn-rich deposits, and 200 ° to 350°C for Cu-Pb-Zn-rich deposits (Fig. 1.87) (Shikazono, 1985b). Homogenization temperatures are not same as the formation temperature. Therefore, we need pressure correction to estimate the formation temperatures from homogenization temperatures. However, the homogenization temperatures are in good agreement with electrum-sphalerite temperatures (Shikazono, 1985d) (Figs. 1.88 and 1.89). Therefore, pressure corrections to homogenization temperatures of fluid inclusions necessary to obtain formation temperatures are relatively small (less than 30°C). Salinities of inclusion fluids from epithermal vein-type deposits clearly indicate that the salinities of inclusion fluids from these types of deposits are distinctly different, that is, 20-2 NaC1 equivalent wt% (base-metal vein-type deposits) and 0-3 wt% (Au-Ag vein-type deposits) (Shikazono, 1985b) (Table 1.13). Salinities of inclusion fluids from Kuroko deposits (0.5-5 wt% NaC1 equivalent concentration) are between these two types of deposits. This kind of difference is observed in epithermal deposits in other countries (Hedenquist and Henley, 1985).
1.4.3.2. Estimate of temperatures from the electrum-sphalerite-pyrite-argentite assemblage As noted already, the Ag content of electrum in equilibrium with argentite and FeS content of sphalerite in equilibrium with pyrite are expressed as a function of fs2 and temperature, we can estimate temperature from Ag content of electrum and FeS content of sphalerite.
Miocene-Pliocene Hydrothermal Ore Deposits
125
Filling Temperature ('C) Au-Ag Dep.
150 [
200 i
250 ]
I
•
Sanru Teine-Takinosawa Todoroki-Chuetsu Todoroki-Shuetsu Chitose-Daikoku Chitose-Daikoku No2 Chitose-Fukujin Yatani-Kanizawa Takatama Sado-Ohdachi Nebazawa Seigoshi Taio Fuke-Honpi Ohkuchi Hishikari Arakawa Kushikino
I
3OO ] ]
•
•
350
I I
I
•
I
•
I
•
¢
I
: •
I
Pb-Zn'Mn Dep. Ohe-Senzai Inakuraishi Toyoha-Tajima Toyoha-Hadma Toyoha-lzumo Yagumo-Ohgid
•
I •
I
6
Cu" Pb oZn Dep. Oppu Osarizawa AnMnari Hosokura-Shoko Hosokura-Ohtake Hosokura-Hakuho Ohizumi Yatani-Honpi Nanetsu Tochigi Ogoya Taishu-Okutomi Taishu-Shinotomi Taishu-Himi Taishu-Taisho Taishu-Misoge Taishu-Tsurue Taishu-Akushidani Taishu-Shirocake Taishu-Amanohara
•
I
¢
¢
•
I
I •
]
I [
•
]
I
Figure 1.87. Summary of filling temperatures of fluid inclusions from Neogene vein-type deposits in Japan. Solid circIe represents average filling temperatures of fluid inclusions for individual deposits (Shikazono, 1985b).
Chapter 1
126 350
300
_~ 250
g
200
J1 ~
'I l
150
100 150
200 250 300 350 Electrum-Sphalerite Temperature (°(3)
400
Figure 1.88. Electrum-sphalerite temperatures vs. homogenization temperatures of fluid inclusions from epithermal Au-Ag vein-type deposits in Japan. Electrum-sphalerite temperatures were calculated from the iron content of sphalerite and the silver content of electrum (Equation 6 in Shikazono (1985a)). The average and range of homogenization temperatures and electrum-sphalerite temperatures for a given ore deposit are represented by a solid circle and line, respectively. 1 = Ohmidani-Fusei, 2 = Todoroki-Chuetsu, 3 = TodorokiShuetsu, 4 = Taio no. 9, 5 = Toyoha-Tajima, 6 = Nawaji, 7 = Seigoshi no. 2, 8 = Toi, 9 = Yugashima, 10 = Yatani-Tengu, 1l = Yatani-Kanizawa, 12 = Oh~Senzai, 13 = Ohe-Senzai, 14 = Sado, 15 = NebazawaManzai no. 3, 16 = Kamioka, 17 = Chitose-Daikoku, 18 = Yatani Honpi, 19 = Yunoura (Shikazono, 1985a).
In order to obtain reliable formation temperatures based on electrum-sphalerite geothermometer, the following conditions have to be satisfied. (1) The coexisting sphalerite, pyrite, electrum, and argentite must have been at e q u i l i b r i u m at the time of their precipitation. Although it is difficult to evaluate this condition, it is c o m m o n l y observed that these minerals are in direct contact with each other without evidence o f mutual replacement texture. Therefore, it is likely that these minerals have been precipitated nearly contemporaneously. (2) The FeS content o f sphalerite and the Ag content of electrum have not c h a n g e d considerably during the post-depositional period. It is unlikely that the sphalerite c o m p o s i t i o n c h a n g e d d u r i n g the cooling stage, because this mineral is one of the most refractory sulfide minerals. In general, the Ag c o n t e n t of electrum with a large grain size increases from core to margin. Such regularity o f compositional z o n i n g observed in a large grain, wide c o m p o s i t i o n a l variation in a large grain, and the relationship between grain size and compositional range in a grain suggest that the electrum composition is retained d u r i n g the post-depositional period. The A u - A g interdiffusion coefficient for electrum has b e e n experimentally studied ( C z a m a n s k e et al., 1973). The values of the A u - A g interdiffusion coefficients, the compositional z o n i n g pattern in electrum, and the temperature range of c o n c e r n (ca. 2 0 0 - 3 0 0 ° C ) suggest that the post-depositional change
M i o c e n e - P l i o c e n e Hydrothermal Ore Deposits
127
-5
-10
Y~
t~
0 _..1
-15
i
v
I
I
t
i
I
200
!
250
i
i
i
Temp.(°C)
I
i
300
Figure 1.89. Activity of S2(as~)-temperature diagram showing possible as2 and temperature ranges for epithermal Au disseminated-type (hot spring type), epithermaI Au-Ag vein-type and epithermaI base metal vein-type deposits in Japan (Shikazono 1986; Shikazono and Shimizu, 1988b). TABLE 1.13 Filling temperature and NaCI eq. concentration of fluid inclusions from epithermaI gold-silver and base-metal vein-type deposits (Shikazono and Shimizu, 1992) Deposit
Filling temprerature (°C)
NaC1 eq. concentration (wt%)
Gold-silver type
Sado Seigoshi Yatani Ohguchi Todoroki Koryu Chitose Kushikino
305-190 243-178 273-209 265-184 24(~ 122 300-140 300-220 250-210
2.5-1.0 2.8~0.0 1.5~).5 1.64).0 1.7-0.4 1.4-0.0 2.0-1.0 1.1~0.6
Base-metal type
Toyoha Oppu Osarizawa Hosokura Nanetsu Taisyu Asahi Ani Oe Jokoku
300-150 330-170 268-156 231 - 130 306-250 376-150 275 250 266-207 310-145 250-125
4.24).2 i8.3-1.7 7.5~0.0 9.1 ~).0 7.9-4.4 33.5~.0 10.3-8.8 12.0 2.0-0.0 6.0-3.0
128
Chapter 1
in electrum composition may be negligible. One of the most likely explanations for the compositional zoning in a grain is that it reflects the change in chemical parameters of ore fluids during the precipitation of electrum. (3) The value of the activity coefficients of FeS in sphalerite determined for temperatures above 300°C can be extrapolated to lower temperatures. As stated by Barton and Toulmin (1966), ?/FeS does not depend on temperature above about 270°C. However, the activity coefficient below 270°C has not been studied. Scott and Kissin (1973) have stated that activity coefficients for FeS in sphalerite at low temperatures may be substantially different from those at higher temperatures. (4) The effects of impurities such as Mn and Cd in sphalerite on the equations are negligible. Generally, the concentrations of minor elements in sphalerite from the epithermal A u - A g vein-type deposits in Japan, except for iron, are very small (Cd less than 1 wt%, Mn less than n 1 x 10 -1 w t % , Cu less than n x 10 -1 wt%, the concentrations of other elements are also less than n x 10 .2 wt%; e.g., Shikazono, 1978b). These low concentrations affect the fs2-temperature relations in the F e - Z n - S system (e.g., Barton and Toulmin, 1966). Impurities in electrum such as copper and antimony are also very small, generally less than 1 wt%. Therefore, it is likely that these elements do not affect the thermochemical properties of electrum. Argentite from epithermal Au-Ag vein-type deposits in Japan contains sometimes up to 10 wt% selenium (Shikazono, 1978b). However, if the selenium contents are small, it is likely that the activity coefficient for Ag2Se in argentite does not deviate significantly from unity, because complete solid solution between Ag2Se and Ag2S exists above ca. 180°C (Sugaki et al., 1982); However, the absolute value of the activity coefficient for Ag2S in argentite cannot be determined. Impurities in pyrite such as nickel and cobalt are also very low (less than n x 10 -1 wt%), and thus, an activity coefficient of FeS2 in pyrite equal to one can be safely assumed. (5) The effect of pressure is negligible. These epithermal A u - A g vein-type deposits have formed in a shallow and low-pressure environment and pressure correction, such that any correction to the homogenization temperatures will be small (probably less than 20°C). The temperatures estimated from the electrum-sphalerite-pyrite-argentite assemblage are plotted versus homogenization temperatures of fluid inclusions as shown in Fig. 1.88. Although the electrum-sphalerite temperatures and homogenization temperatures from given veins for each deposit have some variation, the averages of the electrum-sphalerite temperatures show a good correlation with the average homogenization temperatures. Most electrum-sphalerite average temperatures correlate to within 30°C of the respective average homogenization temperatures. This good correlation can be seen in Fig. 1.88. The good correlation between homogenization temperatures and electrumsphalerite temperatures suggests several points: (1) the uncertainties of electrumsphalerite temperatures are less than 20 ° to 30°C, even at temperatures from ca. 180 ° to 300°C, (2) the electrum-sphalerite-pyrite-argentite assemblage was formed close to equilibrium in Japanese epithermal Au-Ag vein-type deposits, and (3) the pressure corrections to homogenization temperatures for Japanese epithermal Au-Ag vein-type deposits is small, less than 20°C to 30°C. t All n in this text book as natural number 1 through 9.
Miocene-Pliocene Hydrothermal Ore Deposits
129
1.4.3.3. Gas fugacities
Sulflirfugacity (fs2)
As will be mentioned in section 2.4.3, i s 2 c a n be estimated based on the Ag content of electrum coexisting with argentite (or acanthite), the FeS content of sphalerite coexisting with pyrite and temperature estimated from homogenization temperatures of fluid inclusions. Figures 1.68 and 1.69 show the FeS content of sphalerite and the Ag content of electrum from epithermal Au-Ag vein-type, epithermal base-metal vein-type and Kuroko deposits, indicating different fs2-temperature regions for these types of deposition. Figure 1.89 shows typical range of fs2 and temperature for epithermal basemetal vein-type and Au-Ag vein-type deposits. It is noteworthy that the ranges of fs2 for epithermal Au-Ag, epithermal Au-bearing base-metal, and epithermal Au-free base-metal vein-type deposits are different, while temperatures are not different. As mentioned already, small amounts of electrum occur in epithermal base-metal vein-type deposits. Electrum is not observed in the epithermal base-metal vein-type deposits in which pyrrhotite occurs (e.g., Toyoha-Soya, Oizumi, and Hosokukura Pb-Zn deposits). However, electrum is found in epithermal base-metal vein-type deposits in which hematite is commonly observed (e.g., Osarizawa and Ani C u - P b - Z n deposits). This indicates that electrum precipitates in relatively high fs2 and f02 condition.
Oxygenfugacity (fo2). The fo2-pH diagrams (Figs. 1.90 and 1.91) were constructed at 200°C and 250°C based on the homogenization temperatures and electrum-sphalerite temperatures (Shikazono, 1985d). The ionic strength and activity coefficients of aqueous species are estimated from the freezing temperature of fluid inclusions. The ionic strength is assumed to be 1. Estimates of the total dissolved sulfur concentrations of ore fluids (ZS) responsible for several veintype deposits in Japan are of the order of 10-2-10 -3 mol/kg H20 (Shikazono, 1974b; Hattori, 1975). Chemical analyses of the hot springs accompanied by epithermal basemetal depositions (White, 1967; Browne and Ellis, 1970; Mayhon and Finlayson, 1972; Weissberg et al., 1979) give the values of 10 2-10-3 mol/kg H20. Therefore, the total dissolved sulfur concentration (ZS) was assumed to be 10-2-10 -3 mol/kg H20 for the construction of the diagrams. Gangue minerals and salinity give constraints on the pH range. The thermochemical stability field of adularia, sericite and kaolinite depends on temperature, ionic strength, pH and potassium ion concentration of the aqueous phase. The potassium ion concentration is estimated from the empirical relation of Na+/K + obtained from analyses of geothermal waters (White, 1965; Ellis, 1969; Fournier and Truesdell, 1973), experimental data on rock-water interactions (e.g., Mottl and Holland, 1978) and assuming that salinity of inclusion fluids is equal to mNa+ + mK+ in which m is molal concentration. From these data potassium ion concentration was assumed to be 0.1 and 0.2 mol/kg H20 for 200°C and 250°C. By giving the values of temperature, ZS, ionic strength, FeS content of sphalerite and Ag content of electrum, we can place a limit on fo2. However, we cannot know whether fo2 lies in the predominance field of reduced sulfur species or that of oxidized sulfur species from the constraints mentioned above.
Chapter 1
130
log fo 2
-35
i I
"".,,..
!
Ka ~, Se IAd -~,~,~_ i i Hm ,.~-r ---4 Mt I
B n,, Py I
~"~:L~."~-,~~ /
..A ~g - .
)
.
,
PY . . . . .-.'.- - - F
:%sit
Po
,
,
l i
I
I
-o"~. ~"
o,
, l
I
|
I
I
I
!
/
I
I
3
4
5
6
?
8
9
10
pH
Figure 1.90. fo2-pH diagram constructed for temperature = 200°C, ionic strength = 1, and ]ES = 0.01 moI/kg H20. Solid lines (1), (2), (3), and (4) are, respectively, (1) 10 tool% FeS in sphalerite; (2) 70 tool% Ag in electrum, (3) 60 mo1% Ag in electrum, and (4) 1 tool% FeS of sphalerite; 0.1 tool% FeS sphalerite is nearly coincident with the Bn + Py/Cp boundary. The dotted and shaded areas A and B represent the possible foz-pH regions for Au-Ag vein-type (Yatani, Seigoshi, Omidani, Asahi) and Pb Zn vein-type (Toyoha, Yatani, Ikuno) deposition, respectively. Ka: kaolinite, Se: sericite, Ad: adularia, Hm: hematite, Mt: magnetite, Bn: bornite, Py: pyrite, Cp: chalcopyrite, Arg: argentite, Nsil: native silver, Po: pyrrhotite (Shikazono, 1978b). The fo2 o f ore fluids responsible for the epithermal base-metal veins might have been in the predominance field o f reduced sulfur species because ( l ) pyrrhotite is occasionally found in these deposits, (2) selenium content of argentite is very low and (3) H2S is dominant in the present-day epithermal base-metal fluids. Implication of selenium content of sulfides will be considered later. Barite is sometimes found in the late-stage of mineralization. Thus, it is likely that fo2 of barite stage lies in the predominance field of oxidized sulfur species. As already discussed, fo2 o f Kuroko ore fluids is considered to lie in the predominance field of reduced sulfur species from the following two reasons; (1) Selenium content of sulfides is very low (Yamamoto, 1974) and (2) H2S is dominant in hydrothermal solution venting from back-arc basins (section 2.3) from which hydrothermal ore deposits being similar to Kuroko deposits form. On the other hand, the ore fluids responsible for epithermal A u - A g vein-type deposits contain appreciable amounts of oxidized sulfur species, together with reduced sulfur species. Oxidized sulfur species/reduced sulfur species ratio is considered to be greater than 1, that is, f o 2 lies in the predominance field of oxidized sulfur species. The reasons for this estimation are: (1) hematite is c o m m o n in the deposits, (2) barite is found
Miocene-Pliocene Hydrothermal Ore Deposits
131
log Io 2 -30
Ka] Se i Ad
~
|
~"'-
'
Hm Mt
~ + ~ - ? - 4 5 - , , , B n + P y "~,
.-'-(3)_ -(2)
CpN ~
--(I) - 4 0 ;__Arg____~' .
.
I PY
.
.
Nsi
.
,
=
\
Po
I
, I I
I
3
I I 11
4
l
\
I
I
I
I
I
f
5
6
7
8
9
10
pH
Figure 1.91. fo2-pH diagram constructed for temperature = 250°C, ionic strength = 1, and lgS = 0.01 mol/kg H20. Is~FeS mole percent lines for sphalerite and the stability relations of some hydrothermaI minerals are also given. The solid lines (1), (2), and (3) are respectively 10, 1, and 0.! mol% FeS in sphalerite; 70 and 60 mol% lines of Ag in electrum are nearly identical to lines (i) and (2). The dotted and shaded areas of A and B represent the possible fo2-pH regions for Au-Ag vein4ype (Yatani, Seigoshi, Omidani, Asahi) and Pb-Zn vein-type (Toyoha, Yatani, Ikuno) deposition, respectively. Bn: bornite, Py: pyrite, Cp: chalcopyrite, Arg: argentite, Nsi: native silver, Hm: hematite, Mr: magnetite, Po: pyrrhotite, Ka: kaolinite, Se: sericite, Ad: adularia (Shikazono, 1978b). in Te-type deposits, although this mineral formed at late stage and not found in Se-type deposits, (3) selenium is usually contained in the deposits, (4) tellurium is concentrated in Te-type deposits, and (5) sulfate ion is abundant, together with H2S in geothermal waters, associated with epithermal A u - A g depositions (section 2.1). Selenium and tellurium contents of sulfides are useful indicators for estimating f Q (Shikazono, 1974a, 1978b). The selenium content o f sulfide is governed by the reaction, (MS) + Se 2 - = (MSe) + S 2 -
(1-37)
where (MS) and (MSe) are MS and MSe components in the sulfide-selenide solid solution M ( S I - x , Sex), respectively. Equilibrium constant o f this reaction (K1-37) is, K1-37 = ( a M S e a s 2 ) / ( a M s a s e 2
).
(1-38)
where a is activity. Therefore, if the activity coefficient ratio, gMSe/gMS, does not deviate from unity and the temperature is constant, the ratio, mMse/mMS, correlates to ase2-~as2- which is represented as functions of f02, pH, temperature, and l ~ S e / 2 S (Shikazono, 1978b).
132
Chapter 1
Log (asga~-)
-1 -2 -3
1
-4
!
-5 -6
I
,
A
I
-7
J
-8 -9
I
I
i
(~)
!(2)
(3)
(4)
I(5) I
I
! i
i
i
11
i
7 2345678 Figure 1.92. Dependence of the ratio aSe2 ~as2
i
|
i
,
,i
i
i
~
,
91011121314
,
pH
pH under the conditions: temperature = 150°C, ionic strength = 1, li~S = 0.01 mol/kg H20, ~2Se = 10-7 mol/kg H20, and ZS = ]~rnr (total reduced sulfur content). (1) H2S-H2Se region, (2) H2S-HSe region, (3) HS HSe region, (4) S2 -HSe- region and (5) S2--Se 2 region (Shikazono, 1978b). on
Figure 1.92 represents the pH dependence of ase2 /as2- in the predominance region of reduced sulfur species. Considering the equation mentioned above, it is thought that this activity ratio correlates well with the selenium content of sulfide, if the effects of the activity of the MSe and MS components in sulfide (M: metal element) are neglected. The activity ratio, ase2 ~as2 , is expected to be relatively high in the sericite region and low in the adularia region. However, the selenium content of acanthite from epithermal A u - A g vein-type deposits in which adularia is a common constituent is higher than that from the epithermal Pb-Zn vein-type deposits in which sericite is the common wall-rock alteration product. Therefore, the difference in the selenium content of acanthite from these deposits cannot be explained by the difference in pH. As the values of the activity ratio, ase2 ~as2 , on the f o 2 - p H diagrams at 150°C (Fig. 1.93) and 300°C (Fig. 1.94) are nearly identical, changes in temperature alone cannot be responsible for the variation in the selenium content, if the temperature dependence of the equilibrium constant, KI-37, is not large. In contrast to the effects of pH and temperature, fo2 has a great effect on the selenium content of sulfides, especially in the predominance region of oxidized sulfur species (Figs. 1.93 and 1.94), where the activity ratio is relatively high and increases rapidly with increasing fo2 at a fixed pH. In the predominance region of reduced sulfur species, the ratio is constant at a fixed pH and has a relatively low value. The wide variation in the selenium content of acanthite from epithermal A u Ag vein-type deposits can be explained by assuming that the acanthite formed within the predominance region of oxidized sulfur species at a constant I2Se/~2S ratio, ionic strength, temperature and pH (Figs. 1.93 and 1.94). The very low selenium content of acanthite from epithermal Pb-Zn vein-type deposits suggests the predominance region of reduced sulfur species (Figs. 1.93 and 1.94).
Miocene-Pliocene Hydrothermal Ore Deposits
133
Io Ifo2 -30
-35 HSO~ (Na, K)SO4
-40 _ ~ -4s
__
~50
H2S
"-'
',HS~ S~-
. . . . . . . . . .
2-
H2SIHS i
i
[
HS S; ;I
i
1 2 3 4 5 6 7 8 9 1011 121314
pH
Figure 1.93. foz-pH diagram with the stability fields of aqueous species in Na-K-H-S Se-O system for the conditions: ]ES = 10 -2 moI/kg H20, ~2Se = 10 7 mol/kg H20, ionic strength = 1, and temperature = 150°C. Dashed lines are the ratio iso-ase2 /as2 in logarithmic units. Stability fields for native sulfur and native selenium and the boundaries between predominance regions of oxidized and reduced selenium species are omitted for clarity (Shikazono, 1978b).
Iogfo2
-20 -25
HSO~ (Na, K)SO~
-30 -35 -40
[
/-ii I--
-45
.
.
_0o
- -,
HS:; f
r
=
i
1 2 3 4 5 6 7 8 9 1011 121314
pH
Figure 1.94. fo2-pH diagram with the stability fields of aqueous species in the N a - K - H - S - S e - O system for the conditions: NS = 10 2 mol/kg H20, ]ESe = 10 7 mol/kg H20, ionic strength = 1, and temperature = 300°C. Dashed lines indicate iso-ase2 ~as2 contours in logarithmic units (Shikazono, 1978b).
Chapter 1
134
Log
a Carb
ZnCO3
1 i
O. -2' -4" -6' -8" -10,
U (2) -5o
-~
(3) t 6g fo 2
Figure 1.95. Activity of component ZnCO3 versus fo2- Carbonate containing ZnCO3 is in equilibrium with sphalerite. Thermochemical calculation was made under the following conditions: temperature = 200°C, ionic strength = 1, E;S = 10 2 m, and pH = 5. (1) CH4 and H2S region. (2) H2CO3 and H2S region. (3) H2CO3 and (Na, K) SO]- region (Shikazono, 1977b). In contrast to the estimation of total dissolved sulfur concentration (ES), the total dissolved selenium concentration (ESe) has not yet been estimated. The Beppu hot springs which are accompanied by A u - A g siliceous sinter, contain about 10 -7 tool/1 Se (Uzumasa, 1965). For the construction of Figs. 1.93 and 1.94, the Se/S ratio of the ore fluids is assumed to be 10 -5 . The difference in selenium content of acanthite from the different types of ore deposits can, of course, also be explained by the difference in Z S e / E S in the ore fluids; i.e., the ore fluids responsible for the formation o f epithermal A u - A g vein-type deposits may have had a higher E S e / E S ratio than that for epithermal P b - Z n vein-type deposits. It is likely that considerable amounts of sulfur were derived from marine rocks (Green tuff) and were incorporated into ore fluids for base-metal veins. The zinc content of siderite was applied to estimate fo2 of ore fluid by Shikazono (1977b). The analytical data on the late-stage siderites coexisting with barite, pyrite and hematite from the Ohmori epithermal A u - A g vein-type deposits that zinc content is high in the range of 0.8-5.8 wt% as ZnO, but the siderite of the early stage of mineralization from the Ohmori and from the Toyoha epithermal P b - Z n - A g vein-type deposits does not contain such a large amount of zinc. The relation between zinc content in carbonate and many physicochemical variables such as f % , temperature, pH, and so on was derived on the basis of the equilibrium between coexisting carbonate and sphalerite (Figs. 1.95 and 1.96). It is predicted theoretically that zinc content increases with increasing fo2 in the oxidized sulfur species and oxidized carbon species region (Fig. 1.95). This estimate is consistent with mineral assemblage containing the siderite. The iron content of sphalerites coexisting with siderite and pyrite from the Toyoha deposits is high (3-12 wt%). On the other hand, sphalerite of later stage Ohmori
Miocene-Pliocene Hydrothermal Ore Deposits
135
Log fo 2 -35 HSO41(Na,K)SO 4
HCO~ CO~-
H2C03 HC03
I
+F;; • p ;..=,~~ . ,.,-:..:. . . . . . . . . . . . ... .t. . . . . . .
;
-40
. . . ~ . : ~ - - -- 2(2)
/
Bn+P, ~ "L ""~; ' ~ ' ' " (3) | ..................... ,,. . . . ~ '.':..( Na, K)S02 Cp H~CO~ ! i I I "~ ~ Z-45 D,,
"I
--
: : I, ' .....
1
"°
-5o
i
i
Ka Se Se:A i
4
~.-.aE~!Bn 2 ',~.-... CO.C ',,1.'>E-..J ~I ;! ~"C HL4 T M Po',M t
H2s HS- Hg S2-
t 2
} ] CH4 : I -"~. :j I i, J._,. -~ ~" '
5
6
7
\
\
! g
lb
pH
Figure 1.96. Log foz~H diagram constructed for temperature = 200°C, ionic strength = 1, }IS 10 -2 m, and P,C = 10 tm. Solid line represents aqueous sulfur and carbon species boundaries which are loci of equal molalities. Dashed lines represent the stability boundaries for some minerals. Ad: adularia, Bn: bornite, Cp: chalcopyrite, Ht: hematite, Ka: kaolinite, Mt: magnetite, Po: pyrrhotite, Py: pyrite, Se: sericite. Heavy dashed lines (1), (2), and (3) are iso-activity lines for ZnCO3 component in carbonate in equilibrium with sphalerite: (1 "~ a carb = 0 . 1 . (2) carb carb -(Shikazono,1977b). J ZnCO3 azncO3= 0 . 0 1 . (3) azncO 3-0.00I =
deposits contains small amounts of iron (0.1-0.4 wt%), and the content is narrow in range (although the iron content in early-stage is higher; 0.4-6.2 wt%). The decrease o f iron content from the early to the later stage indicates that oxidation occurred a n d / o r temperature decreased during the formation of the Ohmori deposits. The difference in iron content for these deposits means that the Toyoha deposits formed under relatively lower fo2 conditions than the early stage Ohmori deposits. These conclusions derived from the analytical data on coexisting siderite and sphalerite are consistent with the studies on the chemical environment of epithermal vein-type deposits in Japan (Shikazono, 1973, 1974a; Hattori, 1975).
(fc02), The fCO2 values can be estimated from (1) gangue mineral assemblages including carbonates and (2) fluid inclusion analyses. Shikazono (1985b) summarized the assemblage and mode of occurrence o f common gangue minerals from more than 70 Neogene epithermal vein-type deposits in Japan. Carbon dioxidefugacity
Chapter 1
136
3
J
2 1
~ 0 o M , , ~ k , , ~_
/i
.,d
-2 -3 -4
150
200
250
300 Temperature (°C)
350
Figure 1.97. log fco2-temperature diagram showing the univariant equilibrium curves for some gangue minerals. A. 2Ca2AI3Si3012(OH) (clinozoisite) + 3 SiO2 (quartz) + 2CACO3 (calcite) + 2 H20 = 3 Ca2Al2 Si3OI0(OH)2 (prehnite) + 2CO2 (Xpis = 0.30, Xpis: mole fraction of pistacite component in epidote). B. Ca6Si60]7(OH)2 (xonotlite) + 6CO2 = 6CACO3 (calcite) + 6SIO2 (quartz) + H20. C. CaCO3 (calcite) q- TiO2 (rutile) + SiO2 (quartz) = CaTiSiO5 (sphene) + CO2. D. MnSiO3 (rhodonite) + CO2 = MnCO3 (rhodochrosite) + SiO2 (quartz). E. 3KA13Si3OIo(OH)2 (K-mica) + 4CACO3 (calcite) + SiO2 (qua~z) = 2Ca2A13Si3OI2(OH) (clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 4- 2H20 (Xpis = 0.3). P. 3KAI3Si3Om(OH)2 (K-mica) + 4CACO3 (calcite) + SiO2 (quartz) = 2Ca2AI3Si3OI2(OH) (clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 Jr- 2H20 (Xpis = 0.25). G. 3FeCO3 (siderite) + (1/2) 02 = Pe304 (magnetite) + 3CO2, C (graphite) + 02 = CO2. H. 3CaMg(CO3)2 (dolomite) + KAISi308 (K-feldspar) 4- H20 = 3CACO3 (calcite) + 3CO2 4- KMg3(AISi3Om)(OH)2 (phlogopite). I. C (graphite) + 02 = CO2, FeS (pyrrbotite) + 1/2S2 = FeS2 (pyrite), 2H2S(aq) + 02 = $2 4- 2H20(I). J. FeCO3 (siderite) + Fe203 (hematite) = Fe304 (magnetite) + CO2. K. CaA12Si4OI2-2H20 (wairakite) + KAISi308 (K-feldspar) + CO2 = CaCO3 (caIcite) + KAI3Si3Oto(OH)2 (K-mica) + 4SIO2 (quartz). L. CaAI2Si4OI22H20 (wairakite) + CO2 = CaCO3 (calcite) + AI2Si2Os(OH)4 (kaolinite) + 2SIO2 (quartz). M. CaAI2Si4OI2-4H20 (laumontite) + CO2 = CaCO3 (calcite) + AI2Si205(OH)4 (kaolinite) + 2 SiO2 (quartz) + 2 H20 (Shikazono, 1985b).
As already mentioned the major gangue minerals vary with different deposit types; quartz, chalcedonic quartz, adularia, calcite, smectite, interstratified mica/smectite, interstratified chlorite/smectite, sericite, zeolites, and kaolinite in Au-Ag deposits, chlorite, quartz, sericite, calcite, rhodochrosite, siderite and (magnetite) 2 in Pb-Zn-rich deposits, chlorite, sericite, siderite, hematite, magnetite and (epidote) in Cu-rich deposits. Based on the gangue mineral assemblage (Fig. 1.97), homogenization temperatures of fluid inclusions, thermochemical calculations (Fig. 1.98), and analytical data on fluid inclusions (Fig. 1.99), typical ranges of fco2 for the Au-Ag, Pb-Zn-Mn and Cu-Pb-Zn vein-type deposits were determined to be 10 .3 to 1 atm (190°C-250°C) l0 -I to 10 atm (200-250°C) and 10- I to 103 atm (200°C-350°C), respectively (Fig. 1.100). The fco2 of Kuroko ore fluids is close to that of Cu-Pb-Zn vein-type deposits. The fco2 values could be estimated based on FeCO3 content of carbonates coexisting with iron minerals (pyrite, hematite, magnetite, pyrrhotite) and minerals containing 2 Mineral in parentheses occurs in small amounts in each deposit type.
Miocene-Pliocene Hydrothermal Ore Deposits
137
3 2
1
# o
0
~-1
._1
-2 -3 -4
150
200
250
300 Temperature (°C)
350
Figure 1.98. Summary of fcQ-temperature ranges for Au-Ag-rich, Pb-Zn Mn-rich, and Cu-Pb-Zn-rich vein-type deposits based on gangue mineral assemblages and fluid inclusion data. Line a-b: Equilibrium among graphite, pyrite, and pyrrhorite. Line c~l: fc02 vs. temperature curve for fluid containing i tool% CO2 with temperature of first boiling of 300°C. Line e-f: "Plagioclase" + CO2 = calcite + "kaolinite". Line g ~ : fco2 temperature relation for geothermal reservoir waters obtained by Arn6rsson (1984). Line i-j: f c Q vs. temperature curve for fluid containing 1 tool% CO2 with a temperature of first boiling of 350°C. Line ~1: 2Ca2A13Si3012(OH) + 3KAISi308 + 4CO2 + 2H20 = 4CACO3 + 3KAI3Si3OI0(OH)2 + 6SIO2; mole fraction of KAI3Si3010(OH)2 in mica = 0.6, mole fraction of 2Ca2Fe3Si30[2(OH) in epidote = 0.2 (Shikazono, 1985b). 31
yA
2! 1 S
04
O o
0
C
P
o
_.1 -2 -3 -4
1;0
2;o
2;0
330 T e m p e r a t u r e (°C)
Figure 1.99. Estimated fco2-temperature ranges from anaytical data on fluid inclusions and homogenization temperatures (Shikazono, 1986). T: Taishu (Pb, Zn), O: Ohizumi (Cu, Pb, Zn), Y: Yatani (Pb, Zn), Os: Osarizawa (Cu, Pb, Zn), H: Hosokura (Pb, Zn), C: Chitose (Au, Ag), S: Seigoshi (Au, Ag).
iron as solid s o l u t i o n ( i r o n c o n t e n t in s p h a l e r i t e ) . S h i k a z o n o ( 1 9 7 4 a ) t h e o r e t i c a l l y d e r i v e d the r e l a t i o n s h i p b e t w e e n iron c o n t e n t o f c a r b o n a t e s c o e x i s t i n g w i t h i r o n m i n e r a l s as f u n c t i o n s o f f c Q , f Q , a n d others. W e c o u l d e s t i m a t e f c Q u s i n g his m e t h o d . B u t n o q u a n t i t a t i v e a p p l i c a t i o n h a s b e e n c a r r i e d out.
Chapter i
138
3 2
1
020 o,-1 -2 -3 -4
150
200
250
300 350 Temperature (°C)
Figure 1.100. Typical fco2-temperature ranges for Au-Ag-rich, Pb-Zn-Mn-rich, and Cu-Pb-Zn-rich veintype deposits estimated from gangue mineral assemblages, homogenization temperatures of fluid inclusions, and thermochemical calculations (Shikazono, 1985b).
Seleniumfugacity (fse2).
If coexisting electrum, sphalerite, pyrite, argentite and galena are in equilibrium, the relationship between the Ag content of electrum, selenium contents of argentite and galena, iron content of sphalerite, temperature, fsz and fse2 can be derived from the equilibrium relations for the chemical reactions given in Table 1.14. Scott and Barnes (1971) have obtained an equation representing FeS content of sphalerite in equilibrium with pyrite as functions of fsz and temperature. By combining equations (1)-(5) in Table 1.14 and the relation between activity coefficient of Ag in electrum, Ag
TABLE 1.14 Chemical reactions and equations representing the equilibrium relations used for drawing Fig. 1.101 Chemical reactions
Equilibrium relations
Temperature range (°C)
4 Ag + 82 (g) = 2 Ag2S (acanthite)
(I)
log f s 2 = ( - 9 7 9 0 . 2 I / T ) + 4.83
41ogaAg
(1)
25-176
4 Ag + $2 (g) = 2 Ag2S (argentite)
(2)
log fs2 = (--9173.95/T) + 3.61 + 2 log aAg2s - 4 log aAg
(2)
I76-804
4 A g + Se2 (g) = 2Ag2Se (naumannite) (3)
Iog fSe2 = (-- 10644.67/T) + 3.12 + 21ogaAg2sc 4IogaAg
(3)
133-727
+ 21ogaAg2s
4PbS + Se2 (g) = 2Pb2Se
(4)
log iS2 = log fSe2 q- 755.68 -- 0.24
FeS + 1/2S2 (g) = FeS2 (pyrite)
(5)
log fs2 = -- 15460/T + i4.32 - 2 log XFcS (5)
- 2 Iog(apbse/apbs)
(4)
25-327 ca. 400-700
T: absolute temperature; fs~: sulfur fugacity; fSe2: selenium fugacity; aAg2S: activity of Ag2S in argentite; aAg2Se: activity of Ag2Se in argentite; aAg: activity of Ag in electrum; apbse: activity of PbSe in galena; apbs: activity of PbS in galena; XFeS: FeS mole fraction of FeS in sphalerite (Shikazono and Takeuchi, 1984).
Miocene-Pliocene Hydrothermal Ore Deposits -10
139
-10
o=
8~
-15
~ 2 1
~:~ -15
o' o
~-"'~'-
~,--- 22
-20 H.T
-25 . . . . . . . 150
!. . . 200 . . . . . . . . . . 250 300 Temperature (°C)
-25
. . . . . . . 150 . . . . . . . . . . 200 ....
250
300
Temperature (°C)
Figure 1.101. Selenium fugacity-temperature diagram, i3: Argentite (or acanthite)-electmm-galena-Se2(g) equilibrium curve for XAg2S (X: mole fraction) = 0.8, Xpbse = 0.05 and XAg = 0.4. 14: Argentite (or acanthite)-electrum-S2(g) equilibrium curve for XAg2Se = 0.2 and XAg = 0.4. 15: Argentite (or acanthite)--electrum-galena-Se2(g) equilibrium curve for XAg2S= 0.8, XpbSe= 0.05, and XAg = 0.8. 16: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg2Se = 0.2 and XAg = 0.6. 17: Sphalerite-pyrite-galena-Se2(g) equilibrium curve for Xr,bSe= 0.05 and X F e S = 0.0l. 18: Sphalerite~yrite-galena-Se2(g) equilibriumcurve for Xpbse = 0.05 and XFeS= 0.02. 19:Sphalerite-pyrite-galena-Se2(g) equiiibriumcurve for Xpbse= 0.02 and XFeS= 0.004. 20:Sphalerite-pyrite-galena-Se2(g) equilibriumcurve for Xpbse = 0.02 and XFeS= 0.007. 21: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg = 0.5 and XAg2Se = 0.5. 22: Argentite (or acanthite)-galena-electrum-Se2(g) equiIibrium curve for Xpbse = 0.02, XAg2S~---0.5 and Xag = 0.5. 23: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg2Se = 0.5 and XAg = 0.7. 24: Argentite (or acanthite)-galena--electrumSe2 (g) equilibrium curve for XAg2S= 0.5, XpbSe= 0.02 and XAg = 0.7. Estimated ranges of temperature and fSe2 for Kushikino and Takatama are shown in this figure as shaded areas. H.T.: homogenizationtemperature of fluid inclusions.
content of electrum and temperature obtained by White et al. (1957), the relationships between temperature, fs2, fse2, Ag content of electrum, FeS content of sphaterite, selenium contents of argentite (or acanthite), and of galena can be derived (Fig. 1.101). From these relations and analytical data on coexisting galena, sphalerite, argentite and electrum, the formation temperature, fs2, and fse2 can be estimated (Shikazono and Takeuchi, 1984) (Fig. 1.101). To derive the relations shown in Fig. 1.101, unity of activity coefficients of FeS2 in pyrite, Ag2S, and Ag2Se in argentite (or acanthite) was assumed. Activity coefficient of FeS2 in pyrite should be very close to unity, because the concentrations of minor elements (e.g., Ni, Co) in pyrite are very low, less than n x 10 . 2 wt%. Bethke and Barton (1971) have suggested in their experimental study on the distribution of selenium between coexisting galena and sphalerite that the P b S - P b S e system behaves as ideal solid solution at least above 600°C. Ag2S in argentite (or acanthite), FeS in sphalerite and Ag in electrum are 5 mol%, ca. 20 mol%, 1.2-2.4 mol%, and 4 6 - 6 0 atomic%, respectively, for the Kushikino deposits and 2 mol%, ca. 50 mol%, 0.4-0.7 mol%, and 5 0 - 7 0 atomic%, respectively for the Takadama deposits. Detailed descriptions on these deposits can be referred to in Kitami (1973), Sukeshita and Uemura (1976), Yamaoka and Nedachi (1978b), Izawa et al. (1981) and Takeuchi and Shikazono (1984). Based on these analytical data on the minerals mentioned above and thermochemical consideration formation temperature, fs2, and fse2 for these deposits are
140
Chapter 1
estimated. Temperature and fs2 estimated on the basis of fs2-temperature diagram is ca. 220-300°C and 10-9-10 -13 atm for the Kushikino and ca. 150-250°C and 10-11-10 -18 atm for the Takadama. Temperature and fse2 estimated on the basis of fSez-temperature diagram (Fig. 1.101) is ca. 200-300°C and 10-12-10 -18 atm for the Kushikino and ca. 150-250°C and 10-14-10 -23 atm for the Takadama. Homogenization temperatures of fluid inclusions in quartz for the electrumsphalerite-pyrite-argentite-galena stage of the Kushikino and Takadama deposits are in the range of ca. 180-250°C (Takeuchi, 1979; Izawa et al., 1981) and ca. 160-240°C (Yamaoka and Nedachi, 1978b; Watanabe, 1979), respectively. Wright et al. (1965) have found from hydrothermal experiments that continuous solid solution exists in the PbS-PbSe system at 300°C. Coleman (1959) has described specimens covering the entire range of solid solution between galena and clausthalite from vanadium-uranium deposits of the Colorado Plateau type. Thus, it seems likely that the departure from ideal solid solution in the PbS-PbSe system is not large in the temperature range considered here (ca. 180-300°C). However, unfortunately, absolute values of VPbS and YPbSe (Y: activity coefficient) in the temperature range considered here cannot be estimated. It is known that continuous Ag2S-Ag2Se solid solution exists in the temperature range concerned (ca. 180-300°C), although this solid solution is unquenchable. Deviation from ideality for AgzS-Ag2Se solid solution is also not studied. Analytical data on galena, argentite (or acanthite), electrum and sphalerite which all coexist with pyrite in the Se-rich Au-Ag vein-type deposits are available from Kushikino and Takadama Au-Ag vein-type deposits (Kawai, 1976; Takeuchi, 1979). PbSe in galena, Ag2Se in argentite (Takeuchi, 1979) and thermochemical calculations are used to estimate temperature and fs2 and Jse2- The temperature estimated from this mineral assemblage for the Kushikino seems to be slightly higher than the homogenization temperature of fluid inclusions. There are several possible reasons for this discrepancy. They are: (1) uncertainties of free energy changes for the reactions in Table 1.14, (2) changes of chemical compositions and phases during the post-depositions stage, (3) not simutataneous precipitations of electrum, sphalerite pyrite, galena, argentite (or acanthite) and quartz which are studied for the chemical composition and fluid inclusions, and (4) deviation from ideality of Ag2S-Ag2Se and PbS-PbSe solid solutions. It is difficult to determine which is the main cause for this discrepancy. In order to solve this, more detailed studies on this assemblage and thermochemical properties of PbS-PbSe and Ag2S-AgzSe solid solutions are required. Although such a discrepancy exists, it was for the Se-rich A u - A g vein-type deposits (Kushikino and Takadama) that (l) the mineral assemblage of sphaleriteelectrum-argentite-galena-pyrite assemblage is a useful indicator of environmental condition (fse~, fs2 and temperature), (2) precipitation temperature for this mineral assemblage is in the range of 150-300°C, and (3) fse2 is lower than fs2. Based on the .[Se2, fs2 and temperature estimated from this mineral assemblage, we can place a limit on the ranges of the other important chemical parameters such as total dissolved selenium and sulfur contents in ore forming solution responsible for the Se-rich A u - A g vein-type deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
141
1.4.3.4. Chemical composition of ore fluids The solubilities of ore metals depend on several variables such as fs2, fo2, pH, salinity, NS and temperature (Barnes and Czamanske, 1967; Shikazono, 1972b; Barnes, 1979). In the earlier sections, we estimated the ranges of these variables. Therefore, it is possible to calculate chemical compositions of ore constituent elements in ore fluids. For instance, Au concentration of ore fluids responsible for epithermal A u - A g deposits could be estimated from the following reaction. Au + HS + H2S + l / 4 02 ----Au(HS) 2 + 1/2 H20
(1-39)
Seward (1973) experimentally determined the solubility of Au due to this complex and equilibrium constant for the above reaction. Figure 1.102 shows the solubility of Au on l o g f o z - p H diagram calculated based on the thermochemical data by Seward (1973). The solubility of gold is high in neutral and near oxidized sulfur species/reduced sulfur species boundary (Fig. 1.102). It is noteworthy that this region corresponds to the f o z - p H region of epithermal A u - A g vein-type depositions (Figs. 1.90 and 1.91) (Shikazono, 1974a; Hattori, 1975). The solubility of pure gold is shown in Fig. 1.102. However, gold does not occur as pure gold in A u - A g deposits, but as electrum. Therefore, the effect of Ag content of electrum on gold solubility must be considered. Shikazono and Shimizu (1987) estimated NAu/NAg ratio in ore fluids responsible for Au-Ag veins based on the following reaction. (Au)e] + AgC12 + 2H2S = (Ag)el + Au(HS) 2 + 2C1- + 2H +
(1-40)
where (AU)el and (Ag)el are the Au and Ag components of electrum, respectively. Au
I HS021 S0z~
30
HEMATITE
250 °C m:~S = 3 x 10-3
I
40 PYRRHOTITE MAGNETITE
45
H2SI HS- \ 2
I
I
4
6 pH
i I
8
Figure 1.102. fo2-PH diagram at 250°C showing the stability fields of the principal sulfur species and solubiiity contours for gold in mg/kg as Au(HS)~-(HenIey,1984).
Chapter 1
142 From the equilibrium relation for equation (1-40), we obtain, aAg/aAu =
2 2 2 (au2smagcl~K1-40)/ (mau(us)~aci_aH+)
(1-41)
where K1-40 is equilibrium constant for the reaction (1-40), and ?/AgCl~/)/Au(rtS)~ is assumed to be unity. The equation is controlled by temperature, acl , aH2S, pH and mAgc12/mAu(HS)~ratio. We can calculate the E A u / E A g ratio in ore fluids responsible for Au-Ag veins, by giving temperature, salinity, pH and H2S concentration. The temperature is estimated based on fluid inclusion studies and electrum-sphalerite-argentite-pyrite assemblage. The NaC1 equivalent concentration of ore fluids is approximated from freezing temperature data on inclusion fluids, though the final melting temperature of fluid inclusion ice is also affected by CO2 concentration in epithermal ore fluids (Hedenquist and Henley, 1985). The pH values are estimated assuming the equilibrium among K-feldspar, K-mica, and quartz; this in turn allows a calculation of the potassium ion activity. The activity of HzS is estimated based on the equation showing the relation between the partial pressure of H2S gas and temperature of active geothermal waters (Giggenbach, 1980; Arn6rsson, 1985). Using the typical value of these variables and XAg (mole fraction of Ag in electrum) = 0.5, which is a typical value for electrum from epithermal Au-Ag vein-type deposits, E A u / E A g is calculated to be about 0.1, which is similar to that of Broadlands geothermal water (New Zealand) which is associated with gold precipitation. The above calculation is based on the assumption that AgC12 is the predominant Ag species in ore fluids. However, it is possible that silver bisulfide complex could contribute significantly to the transportation of Ag (Henley, 1985; Brown, 1986). If gold and silver bisulfide complexes are the dominant gold and silver aqueous species in Broadlands geothermal water and ore fluids responsible for Japanese epithermal Au-Ag vein-type deposits, the Ag/Au ratio of ore fluids responsible for Japanese Au-Ag epithermal vein-type deposits may be similar to that of Broadlands geothermal water which is about 0.1. tf mci- and pH are assumed to be 1-5 molal and lower than that for K-feldspar-Kmica-quartz equilibrium, respectively, E A u / E A g is estimated to be considerably lower than 0.1. Therefore, E A u / E A g of ore fluids for epithermal base-metal vein-type deposits is thought to be considerably lower than 0.1. The concentrations of base-metals (Cu, Fe, Pb, and Zn) in hydrothermal solution in equilibrium with sulfides (chalcopyrite, pyrite, galena and sphalerite) depend on several variables such as pH, mcl- concentration, temperature, mN2S, and fo2. The relation between the concentrations and these variables can be derived based on the chemical equilibrium for the following reactions. CuFeS2 + 2 C1- ÷ 2 H + + 1/2 02
~---FeS2 +
CuC12 ÷ H20
(1-42)
FeS2 + 2 C1- + 2 H + + H20 = FeCI2 + 2 H2S + 1/2 02
(1-43)
ZnS + 2C1- + 2H + = ZnCI2 + H2S
(1-44)
The ranges of these variables of epithermal ore fluids were estimated in section 1.4.3. Although the ranges of these variables are wide, we could estimate the concentrations of base metal elements if we took typical ranges of these variables.
Miocene-Pliocene Hydrothermal Ore Deposits
143
1.4.3.5. Stable isotopes 6D and 3180. 3D and 3180 of the ore fluids responsible for epithermal A u - A g and basemetal vein-type deposits in Japan have been estimated from analyses of fluid inclusions (Hattori and Sakai, 1979) and minerals (Watanabe et al., 1976). These data are shown in Fig. 1.103. 3D values of ore fluids for epithermal A u - A g vein-type deposits are similar to those of present-day meteoric water values. 3D values of epithermal ore fluids for base-metal vein-type deposits are slightly higher than those of epithermal A u - A g vein-type deposits. This may be due to the boiling of epithermal base-metal ore fluids and involvement of seawater. 3180 values of ore fluids are higher than those of meteoric water values. This is considered to be due to oxygen shift, which was caused by meteoric water-rock interaction. 3180 of ore fluids responsible for epithermal base-metal ore fluids is higher than that for epithermal A u - A g ore fluids. This is due to high extent of water-rock interaction and probably involvement of seawater and igneous water in the ore fluids. In individual deposit, 3180 of minerals varies widely. For example, 3180 of quartz and adularia increases with the stage of mineralization, and correlates to adularia/quartz ratio (A/Q), and Au and Ag grades of ore (Fig. 1.104). Such increase in 8180 is found also in carbonates in the Seigoshi A u - A g deposits and quartz and adularia in the Hishikari A u - A g deposits (Shikazono, 1988a; Shikazono and Nagayama, 1993). This increase is interpreted in terms of boiling of fluids at late-stage (Shikazono, 1989), decrease in water/rock ratio (Shikazono and Nagayama, 1993) and/or an influence of sedimentary rocks (Shikazono, 1999b).
~
.
-20-40t~
J
SG
S.W.
YUG
-60-80-
-100
I
-10
1
I
I
-8
-6
-4
I
I
-2
0
1
2
4
Figure 1.103.3D and 3180 of ore fluids responsibIe for epithermai Au-Ag vein-type deposits in Japan (Hattori and Sakai, I979: Imai et aI., 1998). S.W.: seawater, M.W. line: meteoric water line, KK: Kushikino, SG: Seigoshi, YUG: Yugashima, TK: Takatama, FUK: Fuke, YN: Yatani, KN: Kanisawa (Yatani), HK: Hishikari.
Chapter 1
144 ©
=
=
0
>
-1-
< 1" 0"
-2
-3. 0.5"
i "
B
-4
Pb
-5
L
O, -2.
-3.
• Quartz-Adularla ratio • 6180 Fluid
© ,.d
•
-4 3-1T I
1 3 6-IT
i 6 [
, 9 6-2T
i
/ 12 6-3T
~Au
I
i I5 64T
J I thickness 18 21 I 5-1T I 5-2T I (cm)
Figure 1.104. The relationship among 8180 (in permil) of fluids, minor element contents, and A/Q (adularia/quartz ratio) in the vein from wail rock side (3-IT) to central part (5-2T) (Ryosen No. 5 vein, Hishikari mine, 85 ml E50) (Shikazono and Nagayama, 1993).
8D and 8180 values for epithermal deposits from other countries are summarized in Fig. 1.105 (Field and Fifarek, 1985). The oxygen shift away from the meteoric water line is always observed, but 8D is similar to meteoric water value, suggesting meteoric water source of epithermal ore fluids. Magmatic contribution to ore fluids has not been found except in some ore fluids responsible for the deposits in the other countries; Tui
Miocene-Pliocene Hydrothermal Ore Deposits
145
8180 %0 -20 0
,
-10
0
i " /; Meteoric Water LinV
-60
Ocean Water iSMOW;
= /
//
&TF
Walrakel/Broadlandsl'.~ LarderelloQ~". . . . . Gov~,ar~/aAW C
Magmatic water 4 [ 1 &GS I•CP
Salton SeaZ.~- . . . . . . . .
-80 -100
+20
/
-20 -40
+10
I II
,/.~
~ao/%~
/
Mt. Lassen(~-----.-~ . . . . . . . II/~T .Uu ATE =11
_~/ oO~/e¢~
/,o ,G ,,~ . ~~/ : ,~q# =~_~--oSteamboat Springs /~-~ / / R&•A I " " CO& / / - " /
-120
6"g"
Geothermal H2' • Surface * Subsurface Hydrothermal H2' • Epithermal
//
-140 ;Y~owstone
/
/_® t____J ~,~ # /..~-
.-e
/ -
// /
CA~/'/
/
-160 I
i
I
f
I
I
I
I
I
Figure 1. i05. Distributions of 3D and 3]80 in various waters, minerals, and hydrothermal fluids of epitherma] deposits (FieId and Fifarek, 1985).
(New Zealand) (Robinson, 1974); Finlandia (Kamilli and Ohmoto, 1977), and possibly Comstock Lode (Taylor, 1973). ~13C and 3180 of carbonates. 313C and 3]80 of carbonates have been obtained by
Osaki (1973), Matsuhisa et al. (1985), Shikazono (1988a, 1989), and Morishita (1993) (Fig. 1.106). Roughly, the data lie between the igneous value (-7%0 to -8%o), and marine carbonate value (-1%o to +4%0). These are similar to the variations in carbonates from the Kuroko mine area. However, 313C and 3180 of carbonates from the epithermal vein-type deposits vary more widely than the carbonates from the Kuroko mine area. Some data plot in the region below the igneous carbonate-marine carbonate mixing line. This suggests that meteoric water was involved in the ore fluids. On the other hand, some data deviate from igneous carbonate-marine carbonate mixing line to higher 3180 region. This higher ~180 values are explained in terms of boiling of ore fluids. High 3180 values are obtained for carbonates of late-stage of mineralization (Shikazono, 1988a). Figure 1.107 shows the frequency of 313C of carbonates from epithermal A u Ag vein-type deposits and that from base-metal vein-type deposits. The carbonates are divided into two types: type A and type B. Type A is characterized by: (1) abundant occurrence in each deposit; (2) coexistence with sulfide minerals; and (3) large grain size. Main carbonate minerals are rhodochrosite and Mn calcite, whereas calcite is the main carbonate mineral for type B. Mn-carbonates of type A occur in P b - Z n - M n vein-type deposits. Type B is characterized by; (1) poor amounts in each deposit; (2) coexistence
146
Chapter 1
25 o o
o
2O
oe
og
o
oo
-to
o
o
o ee
o
o oo
•
o
.
o
• o
o g
g15
E O .IQ
e~
•
~dl0
o•
oS t ~ . ~ , , o ~ . " , % , " Oo°o o
-,..;:
o
~5
°o•
°o.**
Oo , *.~
t..~
o
o
o .
•
°o
o
o
ee
o
oe
| ~o
-5 $13C of C a r b o n a t e s ( ' / . , ) Figure 1.]06. 5180-513C of carbonates from Neogene vein-type deposits in Japan (open circle = calcite; solid circle = rhodochrosite and Mn-ca]cite; solid triangle = dolomite; cross = siderite) (Shikazono, ]989).
with late-stage quartz occurring in vugs; and (3) small grain size. Type B carbonate occurs in Au-Ag vein-type deposits and in Cu-Pb-Zn vein-type deposits but not in Pb-Zn-Mn vein-type deposits. The range of 313C of Japanese epithermal veins is similar to that of the other countries which is in a range of -10%o to 0%0 (Fig. 1.108 (Field and Fifarek, 1985). 313C and 3180 of carbonates from southern Kyushu (Hokusatsu gold district) have been studied in detail (Matsuhisa et al., 1985; Morishita, 1993). Morishita (1993) found that the 313C values of hydrothermal solution in the district during the mineralization stages were low (-11%o), compared with that of average crustal carbon (-7%o), suggesting that 313C of hydrothermal solution is controlled by organic carbon in widely distributed sedimentay rocks of the Cretaceous Shimanto Supergroup basement. 313C and 3]80 of carbonates in the Seigoshi mine district, Izu Peninsula, middle part of Honshu were determined by Shikazono (1988a). He showed that early-stage carbonates have 313C and 3180 values of -2.9%0 to +0.6%0 and +1.7%o to +10.2%o, respectively, suggesting a contribution of marine carbonate in Miocene marine sediments (Green tuff), but late-stage carbonates have high 3180 and low 313C, suggesting the effect
Miocene-Pliocene Hydrothermal Ore Deposits
147
25
20
z
5
1 -10
-8
-6
-4
-2
a'~c (%o)
(b)
25
> , 20 E
<
15
0
-Q 10
E z
5
8130
(%0)
Figure 1.107. Frequency histogram for ~ 13C of type-A(a) and type B (b) carbonates (Shikazono, 1989).
of boiling. Osaki (1973) suggested that the higher 3~3C (-2.5%0 to -7.5%0) and ~]80 values (+7.0%o to + 15.0%o) of rhodochrosites from the Jokoku Mn vein in Southern Hokkaido are due to contamination with limestone. In addition to boiling and origin of carbon, temperature and chemical states of dissolved carbon influence 313C of carbonates (Matsuhisa et al., 1985). These detailed studies on individual mine district suggest that carbon in carbonates was derived from the country rocks underlying the ore deposits and oxygen in ore fluids is controlled by origin of ore fluids (mostly meteoric water) and boiling of ore fluids.
~34S of sulfides. A large number of ~34Sdata on sulfides from epithermal base-metal and A u - A g vein-type deposits are available (e.g., Shikazono, 1987b). The 334S data are summarized in Fig. 1.109. The ~34S values of A u - A g deposits range from -7.5%0 to +5%o. However, majority of the ~34S values fall in a narrow range from -1%o to +3%e. ~34S
148
Chapter 1 Geothermal Systems Geysers:
002 HCO 3 whole rock calcite
a'~c (%0) -20
-10
I
I
0
10
I
Salton Sea/Cerro Prieto: Carbonate host altered carb. host carbonates
Broadlands/Wairakei: calcite
Sediment-Hosted Cortez: carbonate host altered carb. host calcite Carlin: carbonate host altered carb. host calcite
g !
Volcanic-Hosted Pueblo Viejo: carbonaceous sed.
Zoned Polymetallic Veins Creede: Sunnyside: Tui: Casapalca:
carbonates carbonates carbonates calcite
Figure 1.108. Distributions of ~13C in epithermaI deposits (Field and Fifarek, 1985).
values of sulfide sulfur from Green tuff-type epithermal Au-Ag deposits are higher than those from Non-Green tuff-type (Shikazono, 1999b) (Fig. 1.110). The other geochemical and mineralogical features (313C and 3180 of carbonates, salinity of inclusion fluids, Ag/Au total production ratio, association of metals, gangue minerals) are also different in Green tuff-type and Non-Green tuff-type of epithermal Au-Ag deposits (Table !.15). These differences can be explained by the influences of country rocks. The Green tuff-type deposits are affected by marine rocks containing marine sulfates and carbonates and interstitial water of seawater origin with high chloride concentration, resulting in higher ~34S, ~13C, salinity, and Ag/Au total production ratio and enrichment of base metals (Cu, Pb, Zn, Mn, Fe), whereas the Non-Green tuff-type deposits are by subaerial rocks. Sedimentary sulfur and organic carbon were involved in ore fluids responsible for the Non-Green tuff-type deposits, resulting in relatively low ~348 of sulfides and 3 I3C of carbonates. 334S values of epithermal base metal deposits are higher than those of the epithermal Au-Ag deposits and range mostly from +3%o to +7%o (Fig. 1.111 ). Although most of 334S values for base-metal deposits lie in this range, ~34S of composite sample of sulfides from the Motokura Cu-Pb-Zn deposits, Ohmori Cu-Ag deposits, Hosokura Pb-Zn deposits, Sasayama Cu-Pb-Zn deposits and Imai-Ishizaki Cu-Pb-Zn deposits are low, that is, +0.1, +1.8, +2.2, -0.9 and -2.1%o, respectively (Shikazono, 1987b; Shikazono and Shimizu, 1993).
Miocene-Pliocene Hydrothermal Ore Deposits
149 [ ] Ginguro Type
if)
Base metal rich
Ct~
t~ e-
[] Type Disseminated [] Type
~[~
<
"6 E z 1
[~ -7
IoIX~lxlX'gqN]X~ I I ~ I,[ 1,1 I,I 1,
N] -6
.5~.~?~NNN -5
-4
-3
-2
-1
0
1
2
3
4
5
6
-l
8~s(%o)
Figure 1.109. Sulfur isotopic compositions of Neogene Au-Ag vein-type and disseminated-type deposits. Sulfur isotopic compositions on the samples from the Yatani deposits (Sample No. YT26 from Zn-Pb vein ~34S = +3.3%0), and HS72050305-YT1, YT24 and NS-3 from Au-Ag vein (average ~34S = -t-3.3%o)) by Shikazono and Shimazaki (1985) are also plotted. "Base-metal rich" implies the sample containing abundant sulfide minerals but no Ag-Au minerals from base-metal rich deposits and also from Ginguro-type deposits (Shikazono, 1987b).
TABLE 1.15 Some characteristic features (~34S, ~I3c, Ag/Au total production ratio, metals produced, gangue mineraIs) of epithermal Au-Ag vein-type deposits of the Green tuff-type and the Non-Green tuff-type in Japan (after Shikazono, 1996)
334S 3~3C Au produced (metric tons) Ag produced (metric tons) Ag/Au (in wt. ratio) Amount of sulfides (Cu, Pb, Zn)
Green tuff-type
Non-Green tuff-type
high (-1%o to +6%o) high (-7%o to 0%o) 136 2586 high (average 19.1) large
low (-7%o to +2%0) low (-12%o to +2%0) 420 1506 low (average 10.7) small
common common common present
abundant rare rare absent
Gangue minerals
Calcite Rhodochrosite Manganoan calcite Barite
These data indicate that (1) ~348 values of sulfides are different in different mine districts; in the region where thick Green tuff occurs (e.g., Hosokura Pb-Zn deposit, ore deposits in Southwest Hokkaido) some values are relatively high, and ~34S values of the ore deposits in the Non-Green tuff region (e.g., Northeast Hokkaido) (Motokura, Khonomai, etc.) are low, (2) ~34S values of sulfides from small base-metal vein-type deposits (e.g., Imai-Ishizaki, Sasayama) are low, but ~34S values of relatively large base-metal deposits (Taishu, Toyoha, Ani, Osarizawa, Ohe, Jokoku) are high. These data suggest that ~34S values of large deposits are affected by reservoir sulfur in deep part, but sulfide sulfur of small ore deposits is influenced by the surrounding rocks having low ~34S values (basement sedimentary rocks).
Chapter 1
150 15
Non-Green tuff-type
>,1o.
0 c-
OLL
5"
0 q 15" Green tuff-type
o>,10c-
O"
E
LI.. 5-
-15
-5
-10
0
5
10
53'S (permil) Figure 1.110. Frequency of ~348 values of sulfides from the Green tuff-type and the Non-Green tuff-type deposits (Shikazono, 1999b).
Cu.Pb.Zn Vein-Type Deposits ¢--
<
"6 $ E Z i
-8
,
-7
a
i
-6
-5
I
-4
t
i
-3 -2
-1
0
2
3
4
5
6
7 8 8~s (%o)
Figure 1.111. Sulfur isotopic values tbr sulfides from base-metal vein-type deposits (Shikazono, 1987b).
There is another explanation for the variations in 334S values of sulfide sulfur. It was cited that oxidation state (Jb2) and pH of ore fluids are important factor controlling ~34S values of ore fluids (e.g., Kajiwara, 1971). According to the sulfur isotopic equilibrium model (Kajiwara, 1971; Ohmoto, 1972), ~34S of sulfides in predominance
Miocene-Pliocene Hydrothermal Ore Deposits
]
151
Pyrrhotite-bearing
-2-1
0 1 2 3 4 5 6 7
8910
~34s (%°)
Hematite-bearing
•
-2-1
j
•
•
•
.
,
,
,
,
0 1 2 3 4 5 6 7
F]
89'16 8~s (%0)
Figure 1.112. Sulfur isotopic compositionof pyrrhotite-bearing(solid) and hematite-bearing (open) samples from base-metal-richdeposits in Greentuffregion (Shikazono, 1987b).
region of oxidized sulfur species is lower than that in predominance region of reduced sulfur species. Figure 1.112 shows that ~34S of sulfides containing pyrrhotite which formed in the predominance region of reduced sulfur species is lower than that containing hematite which formed in predominance region of oxidized sulfur species; ~348 values for pyrrhotite-bearing samples are less than +4.5%o, while those for hematite-bearing samples are more than +4.5 %o (Fig. 1.112). This is not consistent with equilibrium model (Kajiwara, 1971; Ohmoto, 1972). The ~34Svariation in individual deposits such as Yatani (Shikazono and Shimazaki, 1985), Hosokura, Toyoha (Kiyosu, 1977a; Hamada and Imai, 2000), Ohe (Kojima and Sugaki, 1989) and Taishu (Kiyosu, 1977b) have been studied. These data indicate that ~348 of sulfides do not vary widely in individual deposits. Figure 1.113 shows the distributions of ~34S in epithermal deposits in other countries summarized by Field and Fifarek (1985). ~34S of sulfides from volcanic-hosted epithermal deposits and zoned polymetallic veins are mostly in a range of -5%o to +5%o, typically 0%o (Creede, Sunnyside and Cotqui). This range is similar to that of Japanese epithermal Au-Ag deposits, but lower than that of base-metal deposits. Different interpretations of origin of sulfide sulfur with 0%o have been proposed. For example, Casadevall and Ohmoto (1977) suggest that the sulfur of sulfides in Sunnyside deposits (U.S.A.) originated from evaporite-bearing sedimentary rocks based on the assumption of equilibrium fractionation between H2S and SO ] . in the dominant region of oxidized sulfur species. Kamilli and Ohmoto (1977) also prefer the sedimentary sulfate source at Colqui (Peru), but suggest that an igneous origin is possible. For the Creede deposits, sulfate-sulfide sulfur isotopic equilibrium is not attained and igneous origin seems more likely. Distributions of major epithermal Au-Ag vein-type deposits are shown in Fig. 1.114. Green-tuff-type deposits are defined as the deposits occurring in the Green tuff region and Non-Green tuff type as those occurring in the regions other than the Green tuff region.
Chapter I
152 (~34s
Geothermal Systems Iceland:
magmaticsulfur SO4 sulfates H2S pyrite
-10 I
0
(%0) 10
I
-20 I ! m
Yellowstone: SO4
sulfur H2S Salton Sea: sulfides BroadIands: sulfides Wairakel: SO4 sulfates H2S sulfides
B
B m
I m
Sediment-Hosted Cortez: Carlin:
barite diagenetic py sulfides barite diagenetic py sulfides
32
Volcanic-Hosted Tolfa:
sulfates sulfides Pueblo Viejo: sulfates sulfur sulfides Goldfield: alunite pyrite Z o n e d Polymetallic Veins San Juan Moutains: Creede: barite sulfides Sunnyside: sulfates sulfides Rico: sulfides Ouray: sulfides Tui: barite sulfides Guanajuato: sulfides in country rock sulfides in volc. sulfides in ore Casapalca: sulfides Finlandia: barite sulfides Western Cascades: sulfides Golden Sunlight: barite sulfides
m i
i
n B
45
_.B
m
. . . . . . . . . .
Figure 1.113. Distributions of ~34S in epithermal deposits (Field and Fifarek, 1985).
Figure 1.114 demonstrates that the deposits in northeastern Hokkaido, central Honshu, Sado Island and Kyushu deposits are Non-Green tuff type and those in southwestern Hokkaido, Northeast Japan (Tohoku), the Izu Peninsula, and San-in are Green tuff-type.
Miocene-Pliocene Hydrothermal Ore Deposits 133 °
o,
153
138 °
200 ~
,
%
Hokka~
~
40 ° ~
SW Honshu
35°
~
]
--
,¢ hu '
~
/ ~lzu peninsula
e tuff region t I (submarine volcanic region) ~ Non Green tuffregion I.. • 4 (subaerial volcanic region) Figure 1.114. Distributions of major Green tuff-type (solid circle) and the Non-Green tuff-type (open circle) Au-Ag vein-typedeposits in Japan (Shikazono, 1999b). ~ , , f " / ) MY Kyushu~ /
The Green tuff-type deposits are sometimes hosted not only by the Green tuff formation but also by the rocks overlying the Green tuff formation, implying they are younger than the Green tuff age (ca. 61-15 Ma). Examples of these deposits (Seigoshi, Yugashima, and Toi) occur in the Izu Peninsula, central Honshu. Base metal-rich deposits including epithermal vein-type and Kuroko deposits occur in the Green tuff, while base metal-rich deposits are few in the Non-Green tuff region. Sedimentary rocks often occur as host rocks, footwall rocks and basement rocks in the Non-Green tuff mine area. For example, in southern Kyushu, the Shimanto Supergroup shale is dominant as basement and a host rock for epithermal A u - A g vein-type deposits (e.g., Hishikari). ~34S values of sulfide sulfur from epithermal A u - A g vein type deposits obtained are summarized in Table 1.16 and Fig. 1.109. ~34S value of sedimentary rocks of the Shimanto Supergroup which hosts the Hishikari deposits, southern Kyushu is -12%o (Ishihara et al., 1986) which is considerably lower than those of the Hishikari deposits (+1%o to +2%0) (Shikazono, unpublished. This suggests that sedimentary sulfide sulfur is one of the sources of the sulfides and probably igneous sulfide sulfur is the dominant source. Morishita (1993) showed based on carbon isotopic composition of carbonates that carbon of carbonates in the gold-bearing quartzvein in southern Kyushu was derived from the Shimanto Supergroup shale. Imai et al. (1998) considered that hydrogen in the ore fluids was derived from the Shimanto Supergroup shale based on 3D (--60%0 to -100%o) of inclusion fluids in quartz and adularia of the Hishikari veins. These isotopic
Chapter 1
154 TABLE 1.16
~348 values of epithermal Au-Ag deposits in Japan (Shikazono, 1999b) Green-tuff-type (average +2.9%~, n = 40)
Non-Green tuff-type (average -1.4%~, n = 49) Mine
~348 (%0)
Metals produced
Kohnomai Ryuo Koudou SW H o k k a i d o Katsuyama
-3.5 -7.1 -0.7 -0.9 -7.0 -2.3 -7.2 -5.8
Au, Ag Au, Ag Au, Ag
- 1.3
Au, Ag Au, Ag
-2.8
Au
Sado
Kasuga Hishikari
Iwato
Sappro Eniwa Todoroki
-0.2 -0.2 -14.6 -7.2 -4.2
Au, Au, Au, Au, Au,
Chitose Teine Date Mutsu Innai
Ag Ag Ag Ag, Cu Ag, As
-1.5 -0.9 +1.6 +0.7
Au, Ag
-3.6 +1.0 -7.4 +0.2 +3.4 +5.6 +5.0 -0.5 +1.6 -2.0 --2.8, -1.5 -0.8, +5.5, +3.2, -0.2, -0. l, -2.3 +1.8, +0.4, +0.6, +1.5 +0.3, +0.2, 0.0, -1.1 -0.5, -0.8, -3.7, -5.4
Au Au Au Au Au Au Au Au Au Au
Au, Ag, (Cu)
Handa Yatani
Takatama
+5.1 +4.5 +3.2 +3.2 +3.0 +2.5 +1.8 +4.3 +5.9
Au, Ag
+4.4 +5.2 +4.4 +2.1 +2.9 +3.5 +4.5 -0.4
Au, Ag
Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag, Cu Au, Ag
Ag Ag Ag Ag Ag Ag Ag Ag Ag Ag
Au, Ag
Au, Ag Au, Ag
Au, Ag
Central Honshu
Ashiyasu
+ 1.9
Au, Ag, Cu
+2.9 +2.0 +3.6 +4.9 +3.8 +2.4 +3.2 +5.8 +4.3 +3.2 +3.2 +3.1 +3.1 + 1.0 +0.4 +0.0
Au, Ag
Izu Peninsula
Ohito
Kyushu
Bajo Asahi Hoshino Taio Ohkuchi Yamagano Yamada Arakawa Kushikino Akeshi
Metals produced
NE Honshu
Sado Island
Takachi
Shizukari
Au, Ag
Central Honshu
Nebazawa Haguro Koei Masutomi Suzukura
~34S (%o)
NE Hokkaido
NE H o k k a i d o
Kami-oumu Hokuryu Sanru
Mine
Yugashima Nawaji Jyoren Omatsu Rendaiji Kawazu Okuyama Suzaki Mochikoshi Seigoshi Toi
Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag, Cu, Mn Au, Ag Cu Au, Ag, Cu Au, Ag Au, Ag Au, Ag Au, Ag
+0.1
-0. I Au, Ag
S W Honshu (San- in)
Takeno
+2.4
Au, Ag
Miocene-Pliocene Hydrothermal Ore Deposits
155
data support the view that a part of sulfide ore sulfur originated from sedimentary sulfur in the Shimanto Supergroup shale, although sedimentary sulfide is probably not a dominant source, and large amounts of igneous sulfide sulfur were involved. 3180 values of the Hishikari ore fluids estimated from 3180 values of quartz and adularia and homogenization temperature of fluid inclusions are -6%o to 0%0 (Shikazono and Nagayama, 1993) which are higher than those for the other Japanese epithermal A u Ag ore fluids (Fig. 1.103). These data are interpreted in terms of maturity of hydrothermal system (Shikazono and Nagayama, 1993), boiling of ore fluids (Nagayama, 1993b), a contribution of magmatic water (Matsuhisa and Aoki, 1994), or hydrothermal fluids whose oxygen isotopes exchanged for those of surrounding rocks at a low water/rock ratio (Shikazono and Nagayama, 1993), and a contribution of sedimentary heavy oxygen (Imai and Uto, 2001). Combined isotopic data (3180, 3D, 313C, 334S) indicate that the interaction of hydrothermal solution with the Shimanto Supergroup shale is likely cause for the higher 3180 of the Hishikari epithermal A u - A g ore fluids, compared with those of the ore fluids at main stage of A u - A g mineralization for the other epithermal ore deposits. Shikazono (1988a) has found that ~180 values of late-stage calcite in epithermal ore fluids are higher (0%0 to +10%o), compared with fluid inclusion data (Hattori and Sakai, 1979). Shikazono (1988a, 1989) considered that boiling of ore fluids is important for causing higher 3180 of late-stage calcite. However, it is also likely that this high value seems due to the interaction of hydrothermal solution with sedimentary rocks with high 3~80 values. 334S values of the Green tuff-type are higher than those of the Non-Green tuff-type and are close to ~34S of the Green tuff region (Kuroko and skarn-type deposits) fall in a narrow range of -t-2%o to +7%o and the sulfide sulfur in these ore deposits is called as the "Green tuff sulfur" by Shimazaki (1985). It is inferred by him that the "Green tuff sulfur" was derived from a deep and homogenous reservoir such as magma. However, two sources (igneous source and seawater sulfate) of the "Green tuff sulfur" are also likely because 334S values of base metal deposits (epithermal base metal vein-type and Kuroko deposits) lie between those of igneous sulfur (0%0) and seawater sulfate (+20%o). Kawahata and Shikazono (1988) calculated 334S of hydrogen sulfide (H2S) in ore fluids for midoceanic ridge deposits as a function of seawater/basalt ratio at constant temperature by giving 334S and sulfur content of original basalt. Following their calculation method, we could reasonably explain 3348 of base metal deposits in the Green tuff region (Shikazono, 1987b. However, 334S of the Green tuff-type A u - A g veins are slightly lower than those of the "Green tuff sulfur", indicating that sedimentary sulfur was involved, and degree of contribution of igneous sulfide sulfur was not large. Shikazono (1987b) pointed out a possibility of igneous origin for sulfide sulfur of epithermal A u - A g ore deposits based on average 3348 values of sulfides from these deposits ( - 1%o to +3%0). Smaller involvement of seawater sulfate into epithermal A u - A g ore fluids for the Green tuff-type than the ore fluids for base metal vein-type and Kuroko deposits in the Green tuff region is possible. The mixing of the "Green tuff sulfur" and sedimentary sulfur is likely cause for the difference in 334S of sulfide sulfur of two types of deposits, considering stable isotopic and geologic characteristics of the Green tuff-type and Non-Green tuff-type epithermal A u - A g veins.
156
Chapter 1
25 du N
20, 0
"~ 15. i •
~, 10-
0 a D o
5-
•i
~
•
Im~ •
•
•
mOj mmn--~ m •
0 o
-15
-10
Non-Grenn tuff type 0 Sado A Fuke Hishikari
.~
613C of Carbonates (%~)
b
Green tuff type • Nawagi • Kawaza-Kakehashi • Yugashima
• •
Yatani Seigoshi
Figure 1.115.3tSO-~t3Cof carbonates fromepithermal vein-typedeposits in Japan (Shikazono, 1999b).
313C values from both types of deposits are different (Fig. 1.115). 313C of the Non-Green tuff-type carbonates are relatively low, suggesting a contribution of organic carbon in sedimentary rocks in the mine areas. Although the ~13C values of the Green tuff-type vary widely, the data lie between igneous value (-7%0 to -8%0) and seawater value (-1%o to -4%o), suggesting a contribution of these two sources. Total metal (Au, Ag, Cu, Pb, Zn, Mn) production during the past, and Ag/Au total production ratio of major epithermal Au-Ag deposits (Green tuff-type and Non-Green tuff-type) are summarized in Table 1.17. Large amounts of Au and Ag have been produced from the Non-Green tufftype (e.g., Hishikari, Sado, Kohnomai, Kushikino). Ag/Au total production ratio of the Non-Green tuff-type (average 10.7) is lower than for the Green tuff-type (average 19.1). These differences can be explained by the HSAB principle by Pearson (1963, 1968). This principle indicates that HS- and H2S are likely to form complexes with the metals enriched in the Non-Green tuff-type (Au, Hg), whereas CI prefers to form complexes with the metals concentrated in the Green tuff-type (Ag, Pb, Mn, Fe, Cu). Occurrence of gangue minerals in both types of deposits is different. For example, Mn minerals (Mn carbonates, Mn silicates) occur abundantly in the Rendaiji, Yugashima, Yatani, and Todoroki epithermal Au-Ag vein-type deposits in the Green tuff region but not in the Non-Green tuff-type. 334S values of barite from these deposits are high (+18%~
Miocene-Pliocene Hydrothermal Ore Deposits
157
TABLE 1.17 Tonnages of gold, silver and the other associated metal, silver/gold ratio, K-Ar ages and host rocks for the Te-type and Se-typc epithermal gold deposits (Shikazono et ah, 1990) Deposit
Au (ton)
Ag (ton)
Ag/Au
Sanru
6.7
40
6.0
Koryu Chitose Yatani
0.76 22.8 1.7
22.2 105 64
29.4 4.7 39
Takadama Nebazawa Seigoshi Omidani Sakoshi-Odomari Kushikino Hishikari
28.8 1.0 13.5 0.3 1.1 55.2 21.7
279.9 65 455 79 9.7 456.2 14.3
Date Kobetsuzawa Teine Chitose Mutsu Osorezan Sado
<1 <1 10.4 22.8 <1
Other metals
Year of production (ton)
K-Ar ages
Host and country rocks
1925-1974
12.4 4- 0.6
1903-1957 I93(>-1974 1870-1974
1.0 4- 0.3 4.7 3.34.0.3
Rhyolitic tuff, tuff breccia, shale Shale Propylite Acidic tuff, shale
9.7 65 34 267 8.8 8.3 0.6
1429-1974 1942-1974 1935-1976 1914-1974 1977-1982 1965-1986 1983-1990
8.4 5.0-5.7 1-3.7 66-68
<1 <1 62.6 105 <1
0.7-2 48.1 6.0 4.7 0.1
1932-1975 1955 1932-1971 1936-1974 1940
5.2:t:0.4
82.9
2404.3
33
Cu: 5400
1601-1988
Kawazu
5.4
272.8
1-5
1915-1959
Okuyama Suzaki Chugu Takeno
<1 2.0 <1 4.6
<1
Cu: 1054 Mn: 15840 Cu: 1250
13.44-0.514.5±0.5 22.1 4-0.7 24.4 4- 0.8 1.44-0.31.5 4- 0.3
<1 90.0
Agawa Kato
<1 <1
<1 <3
lriki Yamada Fuke Okuchi Kushikino
<1 <1 2.4 21.2 55.2
<1 <7 1.5 I5.7 456.2
Se-type
Cu: 70 Cu: 1270 Pb: 29210 Zn: 58420
4.0 4- 0.3 0.97 4- 0.041.54.0.3
Tuff, shale Rhyolite Andesite, diorite porphyry Slate, sandstone Dacitic tuff Andes±re, propylite Shale, andesite
Te-type
Cu: 7560.8 Cu: 70
0.1 20
10-I2 I054 0.7 0.7 8.3
Cu: 200 Pb: 500 Zn: 100 Cu: 1.3
1912-1963 1914-1941 1938 1868-1949
4.7
17.9 4- 4.518.24-4.6
1933-1945 1933-1954
1933-1937 1916-1955 1896-1947 1905-1974 1865-I 986
Tuff Propylite Propylite Propylite Liparitic tuff Dacite Dacite, rhyolite shale, dacitic tuff
Propylite, rhyolite Propylite, dacite, basalt Propylite, rhyolite Tuff Granite, tuff, andesite Rhyolite, dacitic tuff Porphyrite
0.453 -t-0.018 1.4 4- 0.2 1.i :t- 0.5 4.0 :k:0.3
Tuft, andesite, liparite Propylite Propylite Andesite, rhyolite Propylite, andesite
158
Chapter 1
to +29%o) (Watanabe and Sakai, 1983; Shikazono et al., t990), suggesting that sulfate sulfur of barite was derived from sulfate in the submarine volcanic and pyroclastic rocks (gypsum, anhydrite, sulfate ion in interstitial water) (Watanabe and Sakai, 1983). Carbonates are found both in the Green tuff-type and the Non-Green tuff-type, but large amounts of carbonates (dominantly calcite) occur in the Non-Green tuff-type such as the Kushikino and Kohnomai epithermal Au-Ag vein-type deposits. ~13C and 8180 of carbonates from epithermal vein-type carbonates are plotted in Fig. 1.106. 313C values from both types of deposits are different, gl3C values of the Non-Green tuff-type carbonates are relatively low, suggesting a contribution of organic carbon in sedimentary rocks in the mine areas. Although the ~13Cvalues of the Green tuff-type vary widely, the data lie between igneous value (-7%o to -8%0) and seawater value (-1%o to +4%o), suggesting a contribution of these two sources.
834S and 8180 of sulfates. Watanabe and Sakai (1983) have analyzed 3348and 3180 of barite, anhydrite and gypsum in the epithermal vein-type deposits. They found that (1) the hydrothermal vein sulfates are characterized by more scattered distributions of 334S values ranging mostly from -t-10%o to .1.20%o and 8180 mostly from 0%0 to +14%o than those of Kuroko sulfates (section 1.3.3) (Fig. 1.43). In Fig. 1.43, 334S and 8180 of sulfates from epithermal vein-type deposits (Watanabe and Sakai, 1983) are plotted. These data show that 334S (mostly from .1.24%o to -1-t-37.8%o)and 3180 of barite (0.1%o to ,1,1,18.7%o)from epithermal Au-Ag-Te vein-type deposits are higher than that of epithermal base-metal vein-type deposits (~34S; .1.16.0%o to +24.6%o, 3180; +2.1%o to +12.1%o). These data could be explained by: the sulfur of barite from epithermal Au-Ag-Te deposits came both from volcanic gas (SO2) and marine sulfate, but that of epithermal base-metal deposits came from marine sulfate and oxidation of H2S. 1.4.3.6. Lead isotopes Lead isotopic data on the epithermal deposits together with Kuroko deposits are plotted in Fig. 1.116 (Sato and Sasaki, 1973; Sato et al., 1973, 1981; Sato, 1975; Sasaki et al., 1982; Sasaki, 1987; Fehn et al., 1983). It is evident that lead isotopic compositions of epithermal vein ores are more scattered than Kuroko ores, although averaged values are similar to the Kuroko ores. This variation seems to be due to the difference in crustal materials underlying the ore deposits; Lead isotopic compositions of different ore deposits which formed at different ages in the same district show the same values (Sasaki, 1974). 1.4.3.7. Rare earth elements (REE) Kato et al. (1990) analyzed carbonates from epithermal Cu, Zn-Pb and Au-Ag vein-type deposits in Japan for rare earth elements (REE) and found that (1) calcite from Cu-type is characterized by high La/Yb, high REE concentration and negative Ce anomaly is small, and (2) calcite from Zn-Pb type is characterized by high La/Yb, low REE and negative Ce anomaly. These data suggest that ore fluids for Cu-type were generated under low water/rock ratio condition or were influenced by magmatic water, while meteoric water component was dominant for the Zn-Pb type.
Miocene-Pliocene Hydrothermal Ore Deposits 39.0
,
,
%
159 ' !
i
: ............. l "CO
io
C, .04
uo
:d
' o i: t°
°"~
:
38.5 L .............
.t
SRM-981 ,. ,tie
[.r,
206/204
t*"
15.7
!
i
!
i
/k.
!
!
!
I
o TA
'O
0
15.6 • o e4 SRM-981
F~"7"';q :oL:z.?.°?...}
n W
H
.d"
15.5
206/204 I
15./,
16.9
,
I
17.0
~
t
18.3
I
18./,
,
I
l&5
I.
187
Figure 1.116. Lead isotopic variation in Japanese Neogene ores. The majority of data fall in a relatively narrow range which is no more than twice the experimental uncertainty indicated by the replicate analyses of NBS-SRM-981 standard (Sasaki et aI., 1982).
The REE characterics of calcite from the A u - A g type are variable. For example, calcites from Sado A u - A g vein, one of the largest A u - A g deposits in Japan have both signatures of meteric water and magmatic (or igneous) contributions. Positive Eu anomaly is only found in calcite containing low REE from A u - A g type (Seigoshi deposit) (Shikazono, unpublished). 1.4.4. Se- a n d T e - t y p e A u - A g d e p o s i t s As shown in Fig. 1.117, Se-type and Te-type epithermal A u - A g vein-type deposits are located in the Cretaceous-Quaternary volcanic terrane of Japan (e.g., northeast and southwest Hokkaido, middle Honshu, south Kyushu). Some Te-type deposits are located in regions similar to the Se-type deposits. Sometimes, Te mineralization is associated with the Se-type deposits, though Te minerals usually do not coexist with Se minerals. However, rarely, Te minerals coexist with Se minerals in the Te-type deposits (e.g., Teine, Suzaki, Kawazu, Iriki) on a polished section scale. For example, coexistence of native Te and Se-bearing tetrahedrite is found at Teine. Generally, Te mineralization occurs at
Chapter 1
160
Teine
I oh,os ( 200km
-Kobe~uzawa Date
Osorezan
Mutsu ."
4o~u
Sado
Chugu
a :no
Ornidani
0 0 Kawazu
1 ~=1
Solgoshi Iriki
okuchl MTL
I
~ . 133°
2
I
138°
~
4
s ['7"1
3 r!'N 615 1
Hlshikari Yamanda Figure 1.117. Map showing the distribution of Se-type and Te-type epithermal gold deposits in Japan. 1: Green-tuff and subaerial volcanic region of Tertiary/Quaternary ages, 2: Main Paleozoic/Mesozoic terrane, 3: Main metamorphic terrane, 4: Te-type deposits, 5: Se-type deposits, 6: Te- and Se-bearing deposits, ISTL: Itoigawa-Shizuoka tectonic line, TTL: Tanakura tectonic line, MTL: Median tectonic line (Sbikazono et al., 1990). Kushikino
~
a higher level than Se mineralization in the same mine district. Aoki (1988) reported that gold precipitation is currently taking place from the Osorezan hot springs, north Honshu, Japan. At Osorezan A u - T e minerals (e.g., krennerite, coloradoite) are found at very shallow levels from the surface but no Se minerals have yet been identified (M. Aoki, personal communication, ] 988). K - A r ages data on adularia and sericite in the veins and altered host rocks indicate that ages of mineralization vary widely, ranging from 1 Ma to 68 Ma and from 1 Ma to 24 Ma for the Se-type and Te-type, respectively (Tables 1.17 and 1.18). The total production of gold, silver and other associated base metals and silver/ gold production ratios from these deposits are summarized in Table 1.17. In addition to gold and silver, lead, zinc and manganese have been produced from some of the Se-type (e.g., Yatani) and copper has been produced from some of the Te-type (e.g., Teine, Kawazu). Total tonnage of production of Au and Ag from the Se-type is greater than
Miocene-Pliocene Hydrothermal Ore Deposits
161
TABLE 1.i8 Summary of geologic, mineralogic and geochemical characteristics of the Te-type and Se-type deposits (Shikazono et al., 1990) Se-type Associated metals Ag/Au Host rocks
Au, Ag, Pb, Zn, Mn More than l0 Sedimentary rocks (dominantly shale), volcanic rocks (dacite and andesite) Form of deposits Vein Age of mineralization Late Cretaceous-Quaternary Opaque minerals Argentite,polybasite, naumannite, aguilarite, Ag-rich electrum, pyrargylite, sphalerite (FeS; 1-5 wt. %), galena, tetrahedrite, chalcopyrite, pyrite Gangue minerals Quartz (fine-large grained), adularia, illite/smectite, chlorite/smectite, calcite, rhodochrosite Homogenization 150-270°C temperatures ~34S of sulfides - 8 to +5% ~348 of sulfates Sulfur activity low Oxygen activity low pH high
Te-type Au, Ag, Cu, Bi, (Hg), (TI) Less than I0 Volcanic rocks (dacite and andesite) Vein, massive Miocene-Present Hessite, petzite, native Te, sylvanite, Au-rich electrum, chaIcopyrite, pyrite, marcasite, tetrahedrite, enargite, bismuthinite, sphalerite (FeS: 1 wt. %), galena Quartz (very fine-grained), barite, illite, kaolinite, adularia 200-300°C - 3 to +7%0 +18 to +29%o intermediate intermediate intermediate
those of the Te-type. A g / A u production ratio (in weight) from the Te-type and Se-type vary widely, generally less than 10 for the Te-type and more than 10 for the Se-type. Host rocks for the Se-type and Te-type epithermal A u - A g deposits are summarized in Table 1.17. In general, the dominant host rocks for both deposits types are intermediate and felsic volcanic rocks. Sedimentary rocks (usually shale) sometimes host the Se-type (e.g., Sanru, Kohryu, Takadama, Ohmidani, Hishikari), but never host the Te-type deposits. All o f the Se-type are vein in form, and most of the Te-type are also vein in form (Table 1.18). However, some Te-type (e.g., Kobetsuzawa, Date, Suzaki, Iriki) are stratabound or massive. Total production of Au from the Te deposits is small. Size (length and width of the vein) of the Se-type vein is generally greater than the Te-type vein. Temperatures of formation can be estimated from homogenization temperatures of fluid inclusion. The typical range o f temperature o f formation for the Se-type and Te-type is 150-270°C and 200-300°C, respectively (Tables 1.18 and 1.19). The pressure correction to obtain true formation temperatures based on homogenization temperatures of fluid inclusions has not been carried out because evidence for boiling is found in the fluid inclusions from some deposits such as Chitose (Yajima, 1979), Nebazawa (Enjoji and Nakayama, 1982); furthermore, homogenization temperatures of fluid inclusions are in good agreement with electrum-sphalerite temperatures (Shikazono, 1985d), indicating that the pressure o f formation is close to the vapor saturation curve. Electrum-sphalerite geothermometer can be used to estimate the temperature of formation for some deposits
Chapter 1
162
TABLE 1.19 Temperatures of formations of the Te-type and Se-type epithermal gold deposits estimated from fluid inclusion homogenization temperatures (Shikazono et al., 1990) Deposit
Homogenization (°C) Range
Average
No. of data
Sanru Koryu Chitose Yatani Takadama Nebazawa Seigoshi Omidani Hishikari Kushikino
200-300 190-340 145-349 209-273 162-240 215-260 !95-243 148-178 90-260 210-250
250 255
419
241 200 250 225 160 195 235
30 100 3600 10 232 100
Te-type deposits Teine Chitose Sado Kawazu Okuyama Takeno Kato Fuke Okuchi Kushikino
180-240 220-260 247-305 206-281 193-284 210-270 190-330 220-280 164-240 260-310
265 236 261 250 250 245 230 280
34 31 53 110 16
Se-type deposits
17
of the Se-type but not for the Te-type, because the assemblage of electrum-argentite (or acanthite)-sphalerite-pyrite is absent in the Te-type (Shikazono et al., 1990). The dominant opaque minerals from the Se-type deposits are Se-bearing A g minerals (argentite, acanthite, polybasite, naumannite, aguilalarite, pyrargyrite), electrum, tetrahedrite, sphalerite, Se-bearing galena, chalcopyrite and pyrite (Table 1.20). The dominant opaque minerals from the Te-type are hessite, native Te, petzite, pyrite, marcasite, enargite and bismuthinite (Table 1.21 ). These Te minerals are not present in the Se-type deposits. Very rarely Se-bearing minerals (tetrahedrite, bismuthinite) occur in the Te-type deposits. Generally, sulfide minerals except for pyrite and marcasite are very poor in amount in the Te-type. On the other hand, sulfide minerals such as argentite, sphalerite and galena are abundant in the Se-type. The chemical compositions of opaque minerals (sphalerite, electrum) are different in two types o f deposits. The FeS content of sphalerite from vein deposits of the Te-type is generally lower than that of the Se-type (Fig. 1.118). However, FeS content of sphalerite from massive deposits of the Te-type (Kobetsuzawa, Suzaki) is high, ranging from 1 to 7 mol%. The Ag content o f electrum from the Se-type is higher than that from the Te-type (Fig. 1.119).
TABLE t.20 Dominant opaque and gangue minerals from the Se-type epitb_ermal gold deposits (Shikazono et al., 1990)
¢5
Deposit
Se-bearing minerals
Opaque minerals
Gangue minerals
Sanru
aguilarite, naumannite, poiybasite, pyrargyrite, stephanite
electrum, miargyrite, chalcopyrite, lahore, arsenopyrite, marcasite, pyrite, sphalerite, stibnite cinnabar
quartz, adalaria, kaolinite, seficite, calcite
Koryu
aguilarite, pearceite, polygasite, proustite, pyrargyrite
electrum, miargyrite, native silver, chalcopyrite, lahore, hematite, magnetite, pyrite, galena, sphalerite
quartz, adularia, johannsenite, chlorite, kaolinite, vermiculite, Mn-calcite
Chitose
aguilarite, argentite, pearceite, polybasite, proustite
electrum, chalcnpyrite, lahore, pyrite, galena, sphalerite
quartz, adularite, chlorite, sericite, calcite
Yatani
argentite
electrum, chalcupyrite, marcasite, pyrite, pyrrhotite, galena, sphalerite
quartz, adularia, chlorite, sericite, rhodochrosite
Takadama
aguilarite, freibergitc, naumannite, pearceite, polybasite, pronstite, pyrargyrffc, stephanite
electrum, sternbergite, chalcopyrite, fahore, marcasite, pyrite, galena, sphalerite
quartz, adularita, kaolinite
Nebazawa
argentite, polybasite, proustite, pyrargyrite
electrum, mirargyrite, native silver, chalcopydte, covellite, fahore, arsenopyrite, marcasite, pyrite, galena, sphalerite
quartz, adularia, sericite, calcite
Seigoshi
argentite
electrum, pearceile, polybasite, pyrargyrite, stephanite, chalcopyrite, fahore, galena, sphalerite
quartz, adularia, inesite, xonotlite, chlorite, mixed layer clay mineral, sericite, calcite, rhodochrosite
Omidani
aguilarite, argentite, naamannite
electrum, native silver, pearceite, polybasite, chalcopyrite, lahore, pyrite, galena, sphalerite
quartz, adularia, chlorite, ankerite, calcite
SakoshiOdomari
aguilarite, argentite, jalpaite, naumarmite, polybasite, pyrargyrite
electrum, mckinstryite, native silver, chalcocite, chalct~pyrite, covellite, fahore, famatinite, native copper, arsenopyrite, hematite, pyrite, pyrrhotite, galena, sphalerite
quartz, kaolinite, sericite
Hishikari
naumannite
electrum, chalcopyrite, marcasite, pyrite, galena, sphalerite, stibnite
quartz, adularia, montmorillonite
Kushikino
aguilarite, argentite, naumannite, polybasite, pyrargyrite, stephanite
electrum, pyrostilpnite, chalcocite, ehalcopyrite, covellite, lahore, marcasite, pyrite, galena, sphalerite
quartz, adularia, mixed layer clay mineral, smectite, calcite, zeolite, truscottite
d
r~ ,~
~"
TABLE 1.21
g
Dominant opaque and gangue minerals from the Te-type epithermal gold deposits (Shikazono et aL, 1990) Deposit
Te-bearing minerals
Date
calaverite, coloradoite, electrum, lahore, pyrite tellurantimony altaite, frohbergite, hessite, native dyscrasite, chalcopyrite, hematite, tellurium, rickardite, sylvanite, marcasite, pyrite, pyrrhotite, tellurantimony galena, sphalerite, realgar, stibnite lhhore, gold-fieldite, hessite, native bismuthinite, lahore, hakite, native electrum, bornite, chalcocite, tellurium, rickardite, stuetzite, tellurium, stuetzite chalcopyrite, emplectite, enargite, sylvanite lahore, luzonite, hematite, marcasite, pyrite, galena, sphalerite, orpiment, realgar, stibnite hessite, petzite electrum, cha]copyrite, lahore, pyrite, galena, sphalerite krennerile, native tellurium chalcopyrite, fahore, arsenopyrite. marcasite, pyrite coloradoite, krennerite marcasite, pyrite, orpiment, realgar hessite electrum, polybasite, pyrite empressite, fahore, gold-fieldite. chalcopyrite, fahore, hemusite, kawazulite, native tellurium, hessite, kawazulite, native stannoidite, hematite, pyrite, poubaite tellurium, poubaite, nckardite, sphalerite stuetzite, sylvanke, tellurobithmuthite, tetradymite altaite, hessite, tetradymite electrumm chalcopyrite, native copper, hematite, pyrite, sphalerite calaverite, hessite, kawazulite, clausthalite, kawazulite, native chalcopyrite, marcasite, pyrite, kostovite, krennerite, native tellurium, poubaite, pyrite sphalerite tellurium, petzite, poubaite. sylvanite hessite electrum, argentite, chalcopyfite, electrum fahore, marcasite, pyrite, galena, sphalerite hessite, petzite
Kobetsuzawa
Teine
Chitose Mutsu Osorezan Sado Kawasu
Okuyama Suzaki
Chugu
Takeno
Se-bearing minerals
Opaque minerals
Gangue minerals quartz, anatase quartz, scricite, zeolite, anatase
quartz, sericite, calcite. rhodochrosite, barite
quartz, adularia, chlorite, sericite, calcite quartz quartz, barite quartz, adularia quartz, sericite, barite, anatase
quartz, chalorite, sericite quartz, barite, anatase
quartz.
quartz, adularia, chlorite, kaolinite, sericite, calcite, dolomite
e~
t"5.
TABLE 1.21 (continued) Dominant opaque and gangue minerals from the Te-type epithermal gold deposits (Shikazono e~ al., 1990) Deposit
Te-bearing minerals
Agawa
calaverite, telluro-bismuthite, tetradymite hessite, sylvanite
Kato lriki
fahore, gold-fieldite, native tellurium
Yamada
hessite, tetradymite
Fuke
hes~ite
Okuchi
hessite
Se-bearing minerals
lahore, famatinite
Opaque minerals
Gangue minerals
electrum, bornite, chalcopyrite, hematite, pyrite, molybdcnite electrum, chalcopyrite, pyrite, galena, sphalerite argentite, chalcopyrite, fabore, famatinim, arsenopyrite, marcasite, pyrite, pyrrhotite electrum, chalcopyrite, pyrite, sphalerite electrum, chalcopyrim, hematite, pyrite, galena, sphalerite electrum, argentite,
quartz, sericite, calcite quartz quartz, kaolinite, anatase
quartz, calcite quartz, adularia, chlorite, kaolinite, smectite quartz, adularia,
Chapter 1
166
In the Se-type gangue minerals comprise quartz, adularia, illite/smectite interstratified mixed layer clay mineral, smectite, calcite, Mn carbonates (manganoan calcite, rhodochrosite), Mn silicates (inesite, johansenite) and Ca silicates (xonotlite, truscottite). In comparison, the Te-type contains fine-grained, chalcedonic quartz, sericite, barite, adularia and chlorite/smectite interstratified mixed layer clay mineral. Carbonates and Mn minerals are very poor in the Te-type and they do not coexist with Te minerals. Carbonates are abundant and barite is absent in the Se-type. Grain size of quartz in the Te-type is very fine, while large quartz crystals are common in the Se-type. Hydrothermal alteration patterns in the Se-type and Te-type deposits are not well defined due to a lack of detailed studies of hydrothermal alteration. However, it is evident that propylitic alteration is widely developed in Se-type gold mine districts (Shikazono, 1988a). For example, chlorite and chlorite/smectite interstratified mixed layer clay mineral are abundant in the country rocks which host the Seigoshi, Fuke, Kushikino and Hishikari deposits (Inome et al., 1981; Shikazono, 1985a; Shikazono et al., 2002). Alteration characterized by the presence of adularia is well developed in some Se-type mine districts (e.g., Seigoshi, Hishikari) (Shikazono, 1985b). Argillic alteration overlies the Kushikino, Seigoshi and Takadama vein deposits (Yagyu, 1954a,b; MITI, 1979). The country rocks in Te-type mine districts suffered argillic and sericite alterations and silicification. For example, such alterations are developed in the Kawazu Te-type gold
90 80
E] Te type deposits (vein) [] Te type deposits (disseminated) [] Se type deposits
70 60
~'50 ~ 4o Ii 30 20 ¸ 10
2
3
4
s
6
7
8
1'0 1112 13t;
is
1617
geS Mol% Figure 1.118. Frequency (number of analyses) of FeS content (mole fraction of FeS) of sphalerite from the Te-type and Se-type deposits (Shikazono et al., i990).
Miocene-Pliocene Hydrothermal Ore Deposits
167
Te Type deposits • n=5 Date Okuyama ~l-~n=10 • n=2 Kushikino Agawa Chugu ,~'H. n=15 Fuke = ~n=69 Okuchi n=10 en=17 Chitose : z = n=6 Takeno == n=l Sado Kato SeType deposits
• n=l ~ n=61
Sakoshi Hishikari Koryu Chitose Sanru Kushikino Takadama Yatani Nebazawa Omidani
n=20~
x =
=
~ n=22
4
=
~ n=8
n=41 = : .~ n=44
I
0
I
I
I
I
I
50
I
I
I N Ag
I
I
100
Figure 1.119. Ag content (atomicfraction of Ag) of electrum from the Te-typeand Se-typedeposits (Shikazono et al., 1990). n: number of analyses.
mine district in Izu Peninsula (Watanabe and Nagai, 1986). Advanced argillic alteration closely related to gold mineralization has not yet been recognized in the Se- and Te-type gold deposits. Sulfur fugacity can be estimated from the homogenization temperatures of fluid inclusions, FeS content of sphalerite coexisting with pyrite, Ag content of electrum coexisting with argentite (or acanthite) and the assemblages of sulfide minerals. Typical ranges of sulfur and oxygen fugacities for the Se-type and Te-type are shown in Fig. 1.120. It is likely that sulfur and oxygen fugacities for the Te-type are higher than those for the Se-type at the same temperature conditions. Probable ranges of oxygen fugacity and pH for the Se-type and Te-type are shown in Fig. 1.121. The pH and oxygen fugacity of ore fluids responsible for the formation of the Se-type appears to be lower and higher, respectively, than for those of ore fluids responsible for the formation of the Te-type (Fig. 1.121). Sulfur isotopic compositions (~34S) of sulfides and sulfate (barite) from the Setype and Te-type are summarized in Fig. 1.122. Almost all ~34S values from the Se-type and Te-type fall in a range from -3%o to +6%o (Fig. 1.122). In general, the ~34S values from the Se-type are similar to those of the Te-type. However, some ~348 values from the Se-type are lower than those from the Te-type. The ~34S values of sulfate (barite) sulfur from the Te-type range from +18%o to +29%0 (Fig. 1.122).
Chapter 1
168 -6
#
!
i
i
t~
COVELLITE DIGENITE
~_\'~
PYRITE+BORNITE
/I
ENARGITE
!
FAMACHtNITE F , \ Te-type ,o.;"~
--
......~
\/
~ /
\
1 mole%FeS in SPHALERITE ['-.t~~',"~'~-~-I" ," • ,it."
0
4%
. _//"
i"" s -tyr, e
7-'- % . . ; % . . d . .14
-38
g:
, -37
-35
-33
i -31
Log ao2 Figure 1.I20. Probable ranges of sulfur and oxygen activities for the Te-type and Se-type deposits. The diagram was constructed mainly based on Heald et al. (1987). Temperature = 250°C (Shikazono et aI., 1990).
The above-mentioned geologic, mineralogic and geochemical characteristics of the Se-type and Te-type are summarized in Table 1.18. These characteristics can be used to reconstruct the structure of the fossil geothermal system responsible for the formation of the Se-type and Te-type epithermal gold mineralizations. Figure 1.123 shows the relative sites of Se-type and Te-type epithermal gold deposition in a fossil geothermal system. The Te-type is considered to form a a site closer to the volcanic centre than the Se-type. SO2 gas derived from magma disproportionates in the presence of water to form sulfate ion and hydrogen ion according to the following reactions (Holland and Malinin, 1979). 4SO2 + 4 H 2 0 --+ 3 H2SO4 q--H2S
(1-45)
H2SO4 --+ 2H + + SO 2-
(1-46)
The sulfate ion generated by these reactions precipitated as barite in some Te-type deposits. This hydrolysis reaction can explain the 334S values of barite and sulfides in the Te-type deposits. Some of the lower 334S values of sulfide sulfur in the Se-type deposits than the Te-type deposits could be due to the incorporation of sedimentary sulfur in the host rocks of the Se-type deposits. However, sulfide sulfur of the igneous source was also involved in the ore fluids responsible for the Se-type. Relatively higher sulfur and oxygen fugacities and lower pH and higher temperature of ore fluids responsible for the Te-type than the Se-type may have been caused by a contribution of magmatic (or igneous) sulfur to the ore fluids responsible for the Te-type. However, it cannot be ruled out that seawater sulfate was involved in the ore fluids responsible for the Te-type, because ~34S of barite
Miocene-Pliocene Hydrothermal Ore Deposits -30
I
W!W
I
I
169
J
J ....
I
l
_
~
~L 7 -~,,~
, "1
....._ ~ / ~
......
,.<&
Te-type
ANGLESITE GALENA
'
~O.,
"V',"<'.#_ TENNANTITE
"
I
--I>zl~
#.,
I
]
-
/t
~N'~'~ .
#'; {/
I
'
l
10m?Ie%FeSInSPHA2ERITE I ~/ --
- , o
-
.
.
.
s
.¢"
,I
0,1%
E
~,~~t '
.
i 2
4
pH
6
8
1
10
Figure 1.121. Probable ranges of oxygen activity and pH for the Te-type and Se-type deposits. The diagram was constructed mainly based on Barton et al. (1977) and Heald et al. (1987). Temperature = 250°C, ICS = 0.02 molal, Salinity = 1 molaI with N a / K (atomc ratio) = 9. Dotted area: Te-type, Hatched area: Se-type (Shikazono et al., 1990).
are close to seawater sulfate value and Te-type deposits tend to occur in Green Tuff regions characterized by thick marine sedimentary and volcanic rocks. Advanced argillic alteration is not found in either the Te-type or Se-type deposits. Sericitization, silicification and argillic alteration are common in the Te-type mine districts. Adularia is a characteristic mineral of adularia-sericite-type (Heald et al., 1987). Thus, the Te-type appears to be intermediate type between the acid-sulfate-type and adularia-sericite-type. The Se-type belongs to low sulfidation-type, essentially adulariasericite-type. Almost all Te-type deposits belong to low sulfidation-type but some have characteristics of the high sulfidation-type. Acid-sulfate, high sulfidation deposits form just above the volcanic centre. However, the Te-type form further from the site of the volcanic centre, though close to the centre than the Se-type. The Se-type may occur in basement (mainly sedimentary rocks) as well as young volcanic rocks. The models for epithermal system by Buchanan (1981), Giles and Nelson (1983), and Berger (1985) emphasize vertical change from base metal-type deposits at the deeper part to precious
Chapter 1
170
(a)
[ ] Te type deposits [ ] Se type deposits
¢o
R
"6 E Z
-8
8
.I.I "1"1 "1-1 [ • [/Ij1 "1 "1" • I . I , I . I A J~J I - V I'1/1"1"1"1" • I-I "I "Pq [2t , I Z I ~ I - I ,, V l t l - I - . I - I t l - J : l " - I . I . I . I - I , I , . I . I , -6 -5" -4 -3 -2 -1 0 1 2 3 4 5 6 7 •
..Q
-7
(b)
BARITE
"6 E
[]
Teine
]
Kawazu
]
osorezan
27
F~. 28 29 ~'S(%o)
y.
, l'~, 18 19
20
. if'l, 21 22
23
. ITIQ~TI 24 25 26
Figure I. 122. Frequency (numberof analyses) of (a) ~348 of sulfide and (b) sulthte (barite) minerals(Shikazono et al., 1990). metal-type deposits at the shallower part. However, the above discussion emphasizes lateral changes of epithermal ore-types in the geothermal system, as shown in Fig. 1.123. It is worth comparing the fossil geothermal system associated with epithermal gold deposition with active geothermal systems. Henley and Ellis (1983) and Henley (1985) showed that sulfate-type hot springs occur at the volcanic centre and upper level of geothermal system, while bicarbonate hot springs occur on the margins of the geothermal system. Lateral flow of acidic hot waters from upper level of active geothermal system of high relief, with significant lateral flow, appears to be consistent with the distribution of the Te-type and Se-type. Carbonates are common in the Se-type, and barite is found in the Te-type. The occurrence of carbonates and barite suggests that ore fluids for the Se-type and Te-type are bicarbonate-rich and sulfate-rich, respectively. 1.4.5. Depositional mechanism and origin of ore fluids
1.4.5.1. Depositional mechanism It is widely accepted that boiling of ore fluids took place in the epithermal A u Ag mineralization system from the fluid inclusion data (Nakayama and Enjoji, 1985;
Miocene-Pliocene Hydrothermal Ore Deposits
Sotfatara "
Propylltlc / Z
171
""
~'
Figure 1.123. Schematic model for the formations of the Te-type and Se-type epithermal gold depositions in the fossil geothermal system. Reference: Henley and Ellis (1983) (Shikazono et al., 1990).
Shikazono, 1985a) and stable isotope data (Shikazono, 1988a, 1989). Generally, f s 2 and fo2 increase in response to boiling (Drummond, 1981). If boiling occurs, H2 gas escapes from ore fluids faster than other gaseous species, resulting to higher f02 and fs2- Thus, it seems likely that Au-rich electrum tends to precipitate under the high fs2 and f02 conditions from the ore fluids from which substantial amounts of vapor released. However, if boiling of ore fluids having high CO2 and CH4 concentrations occurs, fs2 and f02 do not considerably increase (Drummond, 1981). This type of boiling or gas loss could lead to the deposition of sulfide minerals (sphalerite, galena, argentite etc.) due to a pH increase (Drummond, 1981; Hedenquist and Henley, 1985). This pH increase is not favorable for the deposition of Au when thio-Au species are dominant among dissolved Au species because solubility of Au due to thio complexes increases with the pH increase (Fig. 1.102). If electrum coprecipitates with argentite under low fs2 and f02 conditions, although the amount of electrum precipitated might be small, considering the solubility of electrum under such conditions, electrum tends to contain high Ag content. Not only the deposition of Au, but also the deposition of Ag has to be taken into account in order to consider the depositional mechanism of electrum. Deposition of Ag from the ore fluids in which Ag chloro complexes (e.g., AgCI~-) are dominant Ag species may be controlled by the following reaction. AgCI~- + 1/2 H20 -+ Ag ÷ 2 C1- + H + ÷ 1/4 02
(1-47)
Thus, a pH increase could be favorable for the deposition of Ag. This may imply that the Ag content of electrum is high when substantial amounts of CO2 loss occur. The CO2 concentration of ore fluids responsible for epithermal base-metal vein-type deposits
Chapter 1
172 NAg
o2
1.0
d
"3 0.8
0.6
23 0
•
16 21 15 • j4 19 ,w
•
•
11 e7 • elO o6 120
*17
o8
25 • 22 200 ~8 e24 26
0.4
5 e4
9 13
27° 0.2
30
,28
r'l
-,
29 Q
;
+,
i
~
i log ( A g / A u )
Figure 1.124. Ag/Au total production ratio from each mine and Ag content of electrum. Solid circle: epithermal Au-Ag vein-type deposits. Open circle: epithermaI base metal vein-type deposits. Solid square: hypo/mesothermal polymetallic vein-type deposits. Open square: epithermal Au disseminated-type deposits. I: Tada, 2: Toyoha, 3: Omidani, 4: Innai, 5: Ikuno, ~" qe-Inakuraishi, 7: Nebazawa, 8: Kawazu, 9: Todoroki, 10: Yatani, 11: Seigoshi, 12: Sado, 13: Takeno, 14.145 ,,awaji, 15: Yugashima, 16: Takadama, 17: Handa, 18: Konomai, 19: Sakoshi-Odomari, 20: Toi, 21: Sanru, 22: Arakawa, 23: Taio, 24: Chitose, 25: Hokuryu, 26: Okuchi, 27: Fuke, 28: Yamagano, 29: Akeshi, 30: Kasuga (Sbikazono, 1986).
in Japan is generally higher than that for epithermal Au-Ag vein-type deposits in Japan (section 1.4.3). It is also inferred that the deposition of argentite occurs by a pH increase, considering the following reaction. 2AgC12 + H2S --+ Ag2S + 4C1- + 2 H +
(1-48)
Therefore, it is likely that Ag-rich electrum and large amounts of sulfide minerals including argentite could precipitate due to CO2 loss and pH increase under low f s 2 and fo2 conditions. Therefore, this mechanism (boiling and gas loss from the hydrothermal solution with different fs2, J%2, CO2 concentration and pH) could explain why the Ag content of electrum correlates with Ag/Au total production ratio (Fig. 1.124). The above mechanism (boiling, loss of CO2 and increase in pH) could also lead to the deposition of other sulfides. The reactions causing sulfide depositions by this mechanism are written as, ZnC12 + H2S -+ ZnS + 2 CI + 2 H +
(1-49)
PbC12 +H2S --+ PbS + 2C1- + 2 H +
(1-50)
CuC1 + 2H2S + FeC12 --+ CuFeS2 + 3 C1- 4- 1/2 H2 + 3 H +
(1-51)
Miocene-Pliocene Hydrothermal Ore Deposits
173
As noted already, Ag content of electrum from the epithermal Au-Ag vein-type deposits is higher than that of the electrum from Kuroko deposits, and Ag/Au total production ratio of epithermal Au-Ag vein-type deposits (average 18) is lower than that of Kuroko deposits (average 76). Therefore, this relation is different from that found in the epithermal vein-type deposits. Ag/Au ratio of electrum may be controlled by the following reaction (Shikazono, 1981): (AU)el + 2H2S + Age12 = (Ag)el + 2 e l - + Au(HS)2 + 2 H +
(1-52)
It is thought from this reaction, that Au-rich electrum precipitates from ore fluids with high C1- concentration and low pH. Therefore, it is considered that different C1concentration and pH are important factors causing different relationship between Ag content of electrum and Ag/Au total production ratio of Kuroko deposits and epithermal vein-type deposits. Temperature is also important factor controlling Ag content of electrum. Shikazono (1981) showed that Ag content of electrum increases with decreasing temperature, assuming that EAg/EAu in ore fluids is constant, and dominant Au and Ag species are Au (HS)2 and AgC12. Figure 1.102 shows the iso-Au concentration contours (solubility of Au) on log fo2-pH diagram (Henley, 1984). The f Q - p H region for maximum Au concentration is close to H2S/HS- boundary and SO2-/reduced sulfur species (HzS, HS-) boundary. The most probable /O2-pH region of epithermal Au-Ag vein formation is also close to SO2-/reduced sulfur species boundary and the pH in equilibrium with adularia/sericite boundary, but the pH lower than adularia/sericite boundary is estimated based on the minerals coexisting with electrum: (1) adularia is abundant in the epithermal Au-Ag vein-type deposits but not in direct contact with electrum; (2) the minerals in direct contact with electrum are quartz, sericite/smectite and chlorite/smectite interstratified clay minerals. The most likely explanation for the variation of f Q - p H region of different epithermal Au deposition is the mixing of ascending hydrothermal solution whose foz-pH region is close to the region showing maximum Au solubility and descending acid sulfate hot spring whose temperature is lower than that of ascending hydrothermal solution. Acid sulfate solution contains appreciable amounts of SO]- but no H2S. This mechanism can explain the formation of Te-bearing Au-Ag veins in which sulfides are poor in amounts. The deposition of sulfides is generally difficult by this mechanism because solubilities of sulfides generally increase with decreasing of pH. However, if temperature of mixed fluid decreases considerably by this mechanism, the deposition of sulfides may be possible, because solubilities of sulfides due to chloro complexes decrease with decreasing of temperature. The above interpretation of formation of epithermal Au-Ag vein-type deposits is supported by (1) thermochemical calculations on this type of mixing (Reed and Spycher, 1985), and (2) the geological occurrence of epithermal Au-Ag vein-type deposits and associated advanced argillic alteration. Reed and Spycher (1985) pointed out based on thermochemical calculations that gold does not precipitate due to boiling because pH increases by boiling, leading to an increase in gold solubility due to gold-thio complexes. They indicated that acidification
Chapter 1
174
,,"0 ~
.,
Y . ~,~ --: [ , 2
I
f~)
J
~
," ',.,, Advanced Argillic ," /~ ,' "~" Alteration , ....,_Funabara area ,,s ~
"Seigoshi
".
-- ,~'~oi ,~l~ ~ ) m i n e , mine ~ ~. ,,1 '/
,"
,"
,"
', J"
"",
,',
F~ "
, ,,.
t
-
,'o
-
Amagi
'~ 5 km
Proov -"
tic
:. mine
\
&
J
: ~,~ Alteration
Mochikoshi
,, Ugusu mine .
~_.J
"
mine -
'
,'
~
Diorite Porphyry Silicified Rock of Advanced Argillic Alteration r - ~ Gold-SilverQuartz Vein Fault
Figure 1.125. Distribution of epithermal Au-Ag vein-type deposits, propyIitic and advanced argillic alterations and intrusive rocks of diorite porphyry (Shikazono, 1985a).
of gold-bearing boiled waters by descending acid-sulfate waters is possibly of great significance to near-surface gold precipitation. The acidification destroys the dominant gold complex, Au(HS) 2, forcing gold precipitation by: H + + Au(HS)2 + 1/2 H2 ~ Au + 2 H2S
(1-53)
Shikazono (1985a) has studied hydrothermal alterations in the epithermal AuAg mine district in Izu Peninsula, middle part of Honshu, and indicated that (1) the propylitic alteration occurs widely in the district; (2) at the centre of the district and stratigraphicalty upper horizon, there exists advanced argillic alteration; (3) epithermal Au-Ag vein-type deposits are distributed at marginal zone in the district (Fig. 1.125); (4) the age of formations of advanced argillic alteration and epithermal Au-Ag veins are nearly the same (Au-Ag vein: 1.8-1.1 Ma; advanced argillic alteration: 2.2-1.2 Ma); and (5) epithermal Au-Ag deposition (Seigoshi) occurred in intermediale f Q - p H region of advanced argillic and propylitic alterations.
Miocene-Pliocene Hydrothermal Ore Deposits
175
By considering the features of hydrothermal system associated with epithermal Au-Ag mineralization, a model for the hydrothermal system is constructed in Fig. 1.123. Schematic diagrams showing the sites of Te-bearing and Se-bearing Au-Ag vein-type deposits in hydrothermal system are constructed based on the mineralogic, geologic and geochemical features of these deposits (Fig. 1.123). SO ] . in the acid sulfate hot springs responsible for advanced argillic alteration and/or epithermal Au-Ag ore fluids is considered to be of volcanic SO2 gas origin because ~348 of sulfates (alunite, barite) in advanced argillic alteration zone and in Te-bearing veins is high, more than +20%0. This is reasonably explained by the generation of SO ] - and H2S by the following hydrolysis reaction of volcanic SO2 gas. 4802 + 4 H20 "-'+ 3 H2SO4 + H2S
(1-54)
It seems likely that the mixing of acid sulfate solution with nearly neutral low salinity Au-bearing fluids seems the most likely mechanism for the formation of epithermal Au-Ag vein-type deposits. This mechanism as a main cause for epithermal-type Au deposition is supported by sulfur isotopic data on sulfides. Shikazono and Shimazaki (1985) determined sulfur isotopic compositions of sulfide minerals from the Zn-Pb and Au-Ag veins of the Yatani deposits which occur in the Green tuff region. The values for Zn-Pb veins and Au-Ag veins are ca. +0.5%0 to +4.5%o and ca. +3%o to +6%0, respectively (Fig. 1.126). This difference in 334S of Zn-Pb veins and Au-Ag veins is difficult to explain by the equilibrium isotopic fractionation between aqueous reduced sulfur species and oxidized sulfur species at the site of ore deposition. The non-equilibrium rapid mixing of H2S-rich fluid (deep fluid) with SOl--rich acid fluid (shallow fluid) is the most likely process for the cause of this difference (Fig. 1.127). This fluids mixing can also explain the higher oxidation state of Au-Ag ore fluid and lower oxidation state of Zn-Pb ore fluid. Deposition of gold occurs by this mechanism but not by oxidation of HzS-rich fluid. This type of mixing could reasonably explain the occurrence of acidic alteration minerals such as kaolinite and alunite in the low-sulfidation epithermal gold vein district (e.g., Seta in northeast Hokkaido, Hishikari in southern Kyushu) (Yajima et al., 1997) Mixing of high temperature hydrothermal solution with high salinity and low temperature solution with low salinity of meteoric water origin seems the most likely mechanism for the base-metal vein-type deposition. The salinity-temperature relationship obtained by fluid inclusion study on the Toyoha and Ohe P b - Z n - M n vein-type deposits and Fujigatani-Kiwada W skarn-type deposits indicates positive correlation (Fig. 1.128). This indicates that the mixing of high temperature fluids of deep-seated origin with low temperature fluid of meteoric water origin, together with the conductive cooling of mixed fluids caused the formation of these ore deposits (Fig. 1.129) (Shibue, 1991). Another possible mechanism for the ore deposition is boiling of fluids. For example, the solubility of galena is controlled by, PbS + 2H + + 2C1- = PbC12 + H2S
(1-55)
If boiling took place, H2S gas removes from the fluids. But CO2 loss causes an increase in pH, leading to the deposition of galena. However, no evidence of boiling of
Chapter 1
176
Galena
[]
Au-Ag
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fluids in inclusion fluids from the base-metal vein-type deposits is observed. Therefore, boiling seems unlikely as a depositional mechanism for galena as well as other common sulfides (chalcopyrite, sphalerite, pyrite).
1.4.5.2. Origin of ore fluids 3D and ~180 data on fluid inclusions and minerals at main stage of epithermal A u A g mineralization clearly indicate that the dominant source of ore fluids is meteoric water. Meteoric water penetrates downwards and is heated by the country rocks a n d / o r intrusive rocks. The heated water interacts with country rocks a n d / o r intrusive rocks and extracts sulfur, Au, A g and other soft cations (e.g., Hg, T1) from these rocks. If hydrothermal solution boils, it becomes neutral or slightly alkaline, leading to the selective leaching of soft cations such as Au, Ag, Hg and T1 from country rocks. However, a contribution of sulfur gas and other components from m a g m a cannot be ruled out.
Miocene-Pliocene Hydrothermal Ore Deposits
177
Au.Ag. ore fluid "~Zn.Pb
(a)
ore fluid
ore fluid
A
~-Zn.Pb
(b) 8 a4Sr.s.s.
ore fluid
D 8 34Sr.s.s.
Figure 1.127. (a) Schematic representation of change in ~SO42-/EH2S ratio (concentrationratio of oxidized sulfur species and reduced sulfur species) and 334S value of reduced sulfur species (334Sr.s.s.) in ore fluids during oxidation of reduced sulfur species at constant temperature and ES (total dissolved sulfur concentration). Dotted area represents the possible region for the fluids accompanied by oxidation of reduced sulfur species. A: initial values of 334Sr.s.s. and zsoZ-/EH2S in fluids prior to the oxidation. B: final values of 834Sr.s.s. and ~SO2 /EH2S in fluids when ~)34Sr.s.s. remains constant during oxidation. C: final values of 834Sr.s.s. and ESO2 /EHzS in fluids when isotopic equilibriumbetween reduced sulfur species and oxidized sulfur species is attained during the oxidation. Hatched areas show 334Sr.s.s. and ZSO]-/EH2S estimated for Zn-Pb ore fluid and Au-Ag ore fluid. (b) Schematic representation of change in ESO]-/EH2S ratio of ore fluids and 334Sr.s.s. value in ore fluids during mixing of t-I2S-richfluid (D) and SO42--rich fluid (E) at constant temperature. Dotted area represents the probable region resulted from the mixing of two fluids (D and E). F: final values of 834Sr~.s. and ESO]-/EH2S of the mixture of C and D fluids when isotopic equilibriumis attained betweenreduced sulfur species and oxidized sulfur species during the mixing. Hatched areas represent ~34Sr.s.s. and ESO]-/EH2S estimated for Zn-Pb ore fluid and Au-Ag ore fluid (Shikazono and Shimazaki, 1985).
8180 of late-stage hydrothermal solution is high (0%o to +3%o), as recognized in the Seigoshi and Hishikari A u - A g veins (Shikazono, 1988a; Shikazono and Nagayama, 1993). This increase in 8180 with the stage of hydrothermal system may be due to the change in water/rock ratio, boiling and kind of rocks interacting with fluids. ~D and 3180 data on fluid inclusions and minerals, 313C of carbonates, salinity of inclusion fluids together with the kind of host rocks indicate that the interaction of meteoric water and evolved seawater with volcanic and sedimentary rocks are important causes for the formation of ore fluids responsible for the base-metal vein-type deposits. High salinity-hydrothermal solution tends to leach hard cations (base metals, Fe, Mn) from the country rocks. Boiling may be also the cause of high salinity of base-metal ore fluids. However, this alone cannot cause very high salinity. Probably the other processes such as ion filtration by clay minerals and dissolution of halite have to be considered, but no detailed studies on these processes have been carried out. Origin of sulfide sulfur of epithermal base-metal veins is thought to be same as that of Kuroko deposits because average ~34S value of base-metal vein-type deposits is +4.7%o which is identical to that of Kuroko deposits (+4.6%o) (Shikazono, 1987b). Namely, sulfide sulfur of base-metal veins came from igneous rocks, sulfate of trapped seawater in marine sedimentary rocks, calcium sulfate (anhydrite, gypsum) and pyrite. ~34S of sulfide sulfur of epithermal base-metal vein-type deposits can be explained by the interaction of seawater (or evolved seawater) with volcanic rocks. There are the following three possibilities for the origin of sulfide sulfur of epithermal A u - A g vein-type deposits:
Chapter 1
178 t-" Q)
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(1) Leaching of sulfide sulfur from subaerial young (Miocene-Pliocene) volcanic rocks (Shikazono, 1987b). (2) Same origin of sulfur for base-metal vein-type deposits and Kuroko deposits. (3) Contributions both from sulfide sulfur leached from volcanic rocks and marine sulfate in Green tuff.
Miocene-t'liocene Hydrothermal Ore Deposits
179
c-
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6
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Figure 1.129. Illustrationof the regressionprocedureof fluid inclusion data on the Ohe deposit(Shibue, 1991).
Possibility (1) was proposed by Shikazono (1987b) who considered that the lower 3348 values of sulfide sulfur than base-metal vein-type deposits and Kuroko deposits can be explained by the leaching of sulfide sulfur from volcanic rocks with lower 334S values (0%0 to +5%o) (Uyeda and Sakai, 1984). (2) is also possible. It has been shown that a large amount of SO ] - together with reduced sulfur species such as H2S and H S - were present in epithermal Au-Ag ore fluids. Therefore, it is likely that 334S of reduced sulfur species whose values are typical Green tuff value decreases due to SO]--H2S(HS - ) fractionation. As noted already, Shikazono (1999b) divided epithermal Au-Ag deposits into Green tuff-type and Non-Green tuff-type based on their distributions. 334S values of Green tuff-type are higher than those of Non-Green tuff-type (Fig. 1.110). Host rocks, basement rocks, distribution of ore deposits, total production of Au and Ag, Ag/Au total production ratio, metals produced, 313C and 3180 of carbonates from the two types of deposits are different (Table 1.18). For instance, 334S values of sulfides from Green tuff-type are higher than those from Non-Green tuff-type. These differences are clearly interpreted by that sulfide sulfur of Green tuff-type was contributed by marine sulfate sulfur but that of Non-Green tuff-type is by igneous and sedimentary sulfide sulfur. The above-mentioned consideration on the origin of Neogene epithermal vein-type and Kuroko deposits is roughly consistent with the view by Mosier et al. (1986). They examined grades, tonnages and basement rocks for 88 epithermal precious and base-metal quartz-adularia-type districts in North and Central America and Japan and revealed that the type of basement rock below the mineralized veins is useful for predicting grade and size of deposits. Epithermal districts overlying basement with salt and evaporites of rocks with trapped sea water have a median tonnage (production and reserves) of 1.4 million metric tons and median grades of 1.5 g/ton Au, 130 g/ton Ag, 2.5% Pb, 1.7% Zn and 0.16% Cu, and districts overlying sedimentay basements have a median tonnage of 0.77 million metric tons and median grades of 7.5 g/ton Au, 110 g/ton Ag, less than 0.025% Zn, less than 0.005% Cu and less than 0.001% Pb, and districts overlying igneous basements have a median tonnage of 0.3 million metric tons and median grades of 5.9 g/ton Au, 38 g/ton Ag, tess than 0.25% Zn, less than 0.002% Cu, and less than
180
Chapter 1
0.003% Pb. Their results clearly demonstrate that basement rocks affect vein components in epithermal precious and base-metal quartz-adularia-type deposits. As noted already, epithermal vein-type deposits are classified primarily on the basis of their major ore-metals (Cu, Pb, Zn, Mn, Au and Ag) into the gold-silver-type and the base-metal-type. Major and accessory ore-metals from major vein-type deposits in Japan were examined in order to assess the possible differences in the metal ratios in these two types of deposits (Shikazono and Shimizu, 1992). Characteristic major ore-metals are Au, Ag, Te, Se and Cu for the A u - A g deposits, and Pb, Zn, Mn, Cu and Ag for the base-metal deposits (Shikazono, 1986). Accessary metals are Cd, Hg, T1, Sb and As for the A u - A g deposits and In, Ga, Bi, As, Sb, W and Sn for the base-metal deposits (Table 1.22, Shikazono and Shimizu, 1992). Minerals containing Cu, Ag, Sb and As are common in both types of deposits. They are thus not included in Table 1.22. As already noted in section 1.4.3, geochemical features of ore fluids responsible for base-metal and gold-silver types of deposits are distinct. They are summarized in Table 1.22. The differences in metals concentrated to the deposits and geochemical fectures of ore fluids responsible for both types of deposits are interpreted in terms of HSAB (hard, soft, acids and bases) principle by Pearson (1963, 1968) below. Ahrland et al. (1958) classified a number of Lewis acids as of (a) or (b) type based on the relative affinities for various ions of the ligand atoms. The sequence of stability of complexes is different for classes (a) and (b). With acceptor metal ions of class (a), the affinities of the halide ions lie in the sequence F - > CI- > Br- > I - , whereas with class (b), the sequence is F - < C1- < B r - < I - . Pearson (1963, 1968) classified acids and bases as hard (class (a)), soft (class (b)) and borderline (Table 1.23). Class (a) acids prefer to link with hard bases, whereas class (b) acids prefer soft bases. Yamada and Tanaka (1975) proposed a softness parameter of metal ions, on the basis of the parameters En (electron donor constant) and H (basicity constant) given by Edwards (1954) (Table 1.24). The softness parameter ~ is given by c~/(oe +/~), where c~ and/~ are constants characteristic of metal ions. They indicated that the softness parameter may reasonably be considered as a quantitative measure of the softness of metal ions and is consistent with the HSAB principle by Pearson (1963, 1968). Wood et al. (1987) have shown experimentally that the relative solubilities of the metals in H20-NaC1-CO2 solutions from 200°C to 350°C are consistent with the HSAB principle; in chloride-poor solutions, the soft ions Au + and Ag + prefer to combine with the soft bisulfide ligand; the borderline ions Fe 2+, Zn 2+, Pb 2+, Sb 3+ and Bi 3+ prefer water, hydroxyl, carbonate or bicarbonate ligands, and the extremely hard Mo 6+ bonds only to the hard anions O H - and 0 2 - . Tables 1.23 and 1.24 show the classification of metals and ligands according to the HSAB principle of Ahrland et al. (1958), Pearson (1963, 1968) (Table 1.23) and softness parameter of Yamada and Tanaka (1975) (Table 1.24). Comparison of Table 1.22 with Tables 1.23 and 1.24 makes it evident that the metals associated with the gold-silver deposits have a relatively soft character, whereas those associated with the base-metal deposits have a relatively hard (or borderline) character. For example, metals that tend to form hard acids (Mn 2+, Ga 3+, In 3+, Fe 3+, Sn 4+, MoO 3+, WO 4+, CO2) and borderline acids (Fe 2+, Zn 2+, Pb 2+, Sb 3+) are enriched in the base-metal deposits, whereas metals that tend to form soft acids
Miocene-Pliocene Hydrothermal Ore Deposits
181
TABLE 1.22 Accessory metals from vein-type deposits in Japan (Shikazono and Shimizu, 1992) Mine
Metal Hg
T1
Cd
Bi
Gold-silver type Konomai Kitanoou Showa Meiji Takadama Hirukodate Okuchi Hazami Osorezan Tsugu Chitose Todoroki Seigoshi Kushikino Teine Nishizawa Kawazu Ifiki Yatani Base-metal type Jizo Akenobe Toyoha Ashio Goka Miyatamata Agenosawa Akarimata Hosokura Nakanosawa Fukoku Ikuno Taishu Suttu Hayakawa Akagane Kishu Kutosan Kurokawa Goka Omidani Imaiishizaki Inakuraishi Tada Fukoku Omodani
Konjo Ryujima
Mo
Sn
W
X X
X
X X
X
X
X
X
X
X
X
X
X
X X X X X X X X X
X
Chapterl
182 TABLE 1.23 Classification of metals and Iigands according to the HSAB principle Mn2+, Ga3+, In3+, Co2+, Fe3+, As3+, Sn4+' MoO3+, WO4+, Co2+ Cu+ , Ag+ , Au+ , T1+, Hg+, Cd2+, Hg2+, Te4+ , TI3+ Fe2+, Co2+, Ni2+, Cu2+, Zn2+, pb2+, Sn4+, Sb3+, Bi3+, SO2 OH , CI-, CO2, SO2 H2S, HS-, S 2
Hard acids: Soft acids: Borderline acids: Hard bases: Soft bases:
TABLE 1.24 Softness parameter of the various metal ions Metal ions
Softness parameter
Metal ions
Softness parameter
Ag+ Hg+ TI+ Cu+ Cd2~ Ni2+ Zn3+ Co2+ Zn2+
1.03 1.01 0.98 0.96 0.96 0.94 0.93 0.92 0.91
CU 2+
0.89 0.87 0.85 0.84 0.82 0.78 0.73 0.58
Bi3+ Pb2+ Fez+ Mn2+ Fe3+ Sn2+ Ga3+
(Ag +, Au +, T1+, Te 4+, T13+) are enriched in the g o l d - s i l v e r deposits. Metals that have high values of the softness parameter (Ag +, Hg +, T1+, Cd 2+) are associated with the g o l d - s i l v e r deposits, whereas those that have low values o f the softness parameter (Zn 2+, In 3+, Bi 3+, Pb 2+, Te 4+, Mn 2+, Sn 4+, Ga 3+) are found with the base-metal deposits. These correlations mean that the HSAB principle could be a useful approach to evaluate the geochemical behavior o f metals and ligands in ore fluids responsible for the formation of the epithermal vein-type deposits. A m o n g the ligands in the ore fluids, H S - and H2S are the most likely to form complexes with the metals concentrated in the gold-silver deposits (e.g., Au, Ag, Cu, Hg, TI, Cd), whereas C1- prefers to form complexes with the metals concentrated in the base-metal deposits (e.g., Pb, Zn, Mn, Fe, Cu, and Sn) (Crerar et al., 1985). Generally, the complexes with intermediate or hard ligands (e.g., chloro complexes) should become more stable with increasing temperature than complexes with soft ligands (e.g., bisulfide complexes) (Seward, 1981 ; Crerar et al., 1985). The higher temperatures of formation of the base-metal deposits (Table 1.13) also are in accordance with the HSAB principle. It was shown in Table 1.13 that the base metal and g o l d - s i l v e r types of deposits formed at different temperatures and concentrations of C I - , H S - , and CO2. Thus, it could further be inferred that the H S A B principle can be successfully applied to the genesis of these vein-type ore deposits, formed mostly at less than ca. 250°C. Crerar et al. (1985) noted that Pearson's rule (the HSAB principle) successfully describes speciation to about 250°C, but may break down at higher temperatures, as all metals become harder.
Miocene-Pliocene Hydrothermal Ore Deposits
183
A few appilications of the HSAB principle to hydrothermal ore deposits have been carried out (Crerar et al., 1985; Wood et al., 1987; Shikazono and Shimizu, 1992). These studies demonstrate that the HSAB principle is useful in interpretations of the metal ratios in ore deposits. For example, Wood et al. (1987) has shown that gold is transported by lower salinity fluids than the base metals, and this difference in salinity is a significant factor in the separation of gold and base metals in Archean deposits in greenstone belts. Cathles (1986) also has indicated that the solubility of gold is much greater in lowsalinity solutions; this can explain the bimodal populations of base-metal-rich, gold-poor deposits (stratiform deposits similar to Kuroko deposits in Japan) and base-metal-poor, and gold-rich deposits in greenstone belts (lode gold deposits). Such differences in the salinity of ore fluids responsible for the epithermal gold deposits and Kuroko and base-metal vein-type deposits in Japan have been pointed out also by Shikazono and Shimizu (1993). Therefore, it could be inferred that the difference in salinity is a main cause for a separation of gold and base-metals in mineralized zones in young (Tertiary to Quaternary) Japanese epithermal vein-type and Kuroko deposits in volcanic terranes and in Archean deposits in greenstone terranes. However, the concentration of CO2 in ore fluids responsible for Japanese gold-silver deposits and Archean lode gold deposits in greenstone belt seems to be different; CO2 concentration is low (0.01-0.1 molal) for Japanese deposits (Shikazono, 1985a), whereas it is high (0.05-2 molal) for Archean deposits (Wood et al., 1987). The reason for this difference is not known.
1.4.6. Hishikari deposit: an example of Japanese epithermal Au-Ag vein-type deposits One of the most important steps made in the research of Japanese epithermal gold deposits during the last two decades is the comprehensive study of the Hishikari gold deposits which occur in southern Kyushu (Fig. 1.130). The deposits are characterized by high gold grade and large amounts of gold reserve. Average gold grade in the Honko (Main) deposits is ca. 70 g/metric ton which is enormously high, compared with the other Japanese epithermal gold deposits. Ore reserve is estimated to be about 250 metric tons Au, which is the largest among the Japanese epithermal gold deposits. Ag/Au production ratio is about 0.7, which is relatively low. The Hishikari deposit is only one which belongs to a giant (Laznicka, 1983) and bonanza (Sillitoe, 1993) epithermal-type deposit (Izawa, 2001). Much research on the Hishikari deposits has been carried out since the discovery of gold veins in 1981. Urashima and Izawa (1983) reported fluid inclusion studies using core samples. Abe et al. (1986) made a detailed description of the veins. The regional geology of this district is described in MMAJ and SMM (1987) and Urashima et al. (1987). Izawa et al. (1990) reviewed the studies on the geology, geophysics and geochemistry which had been done during 1980s. The Special Issue of Resource Geology on the Hishikari deposits (Shikazono et al., 1993) includes various aspects of the Hishikari deposits (oxygen isotopes of gangue minerals, hydrothermal alteration, precipitation sequence, fluid inclusions, vertical electric profiling and electric sounding surveys, structural geological analysis, opaque minerals,
Chapter 1
184
(/-f / -~ k~
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• Hot springtype O EpithermalAu-Agvein-type A Activevolcano
.~v-/
o
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r
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I
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I(km)
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131E Figure ].130. A map showing the location of the Hishikati mine.
prospecting, etc.) which had been studied during late 1980s and early 1990s. No review on the studies during the last decade has been published in English. Thus, a review of the studies done during the last decade and recent studies by the author will be summarized below.
1.4.6.1. Geology and vein system A detailed general geology of the Hishikari district has been done by many investigators (Abe et al., 1986; MMAJ and SMM, 1987; NEDO, 1991; Izawa et al., 1993). The general geology of this district is briefly described below. The district is composed of sedimentary rocks of the pre-Paleogene Shimanto Supergroup (dominantly shale and sandstone) and Quaternary andesitic and dacitic volcanic rocks. The Shimanto Supergroup is comprised of shale, sandstone and their alternations. Although no fossil data are available, the age of sedimentation is thought to be middle to upper Cretaceous age from its lithology (Izawa et al., 1990). The Shimanto
185
Miocene-Pliocene Hydrothermal Ore Deposits
Supergroup rocks in the Hishikari district suffered hydrothermal alteration. Chlorite, quartz and sericite occur abundantly near the veins. The other constituents are pyrite, albite, calcite and organic matter. Quaternary volcanic rocks unconformably overlie the Shimanto Supergroup. Quaternary volcanic rocks consist of Hishikari Lower Andesites (0.98-1.62 Ma), Hishikari Middle Andesites (0.78~0.79 Ma), Kurozonsan Dacites (0.95-1.56 Ma), Shishimano Dacites (0.66-1.6 Ma), and Hannyaji Welded Tuff (0.56-0.731 Ma) (Izawa et al., 1990, 1993). The Hishikari Lower Andesites consist of hypersthene-augite andesite lava flows in the upper horizon and andesitic pyroclastic rocks in the lower horizon. The Kurozonsan Dacites overlying the Hishikari Lower Andesites dominantly consist of hyperstheneaugite-bearing biotite-hornblende dacide lava flows. The Hishikari Middle Andesites overlying the Hishikari Lower Andesites consist of hypersthene-augite andesite lava flows and pyroclastic rocks. The Shishimano Dacites overlying the Hishikari Middle Andesites consist mainly of biotite-hornblende dacite lavas. The Hishikari Upper Andesites overlying the Hishikari Lower Andesites and the Shishimano Dacites consist of hypersthene-augite andesite lava flows and their pyroclastic rocks. The deposit consists of three vein systems, the Honko, Sanjin and Yamada veins (Fig. 1.131). The veins strike N30°E to N50°E and dip 70-90°NW. The veins are hosted
0
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Figure 1.131. Plan of alteration mineral zonation of the volcanic rocks in the Hishikari area (Izawa et al., 1990).
Chapter 1
186
in hydrothermally altered Shimanto Supergroup (Honko (Main) and Sanjin veins), and Quaternary andesite (Yamada veins).
1.4.6.2. Hydrothermal alteration Izawa et al. (1990) recognized the following alteration zones from the vein towards margin of the Hishikari A u - A g mine area, chlorite-sericite zone (zone IV), interstratified clay mineral zone (zone III), quartz-smectite zone (zone II) and cristobalite-smectite zone (zone I) and least altered zone (L.A. (least altered) zone) (Fig. 1.131).
1.4.6.3. Mineralogy Quartz and adularia are dominant gangue minerals in the veins. Small amounts of sericite/smectite, calcite, and truscottite occur as vein staff. Fine-grained electrum, naumannite, Ag-sulfosalts (pyrargyrite, polybasite), chalcopyrite, galena, pyrite, marcasite and stibnite are found. Detailed mineralogical descriptions are given in Urashima and Izawa (1983), Urashima et al. (1987), Izawa et al. (1990), and Nagayama (1993a). The sequence of mineralization has been studied by Nagayama (1993a) and Takahashi et al. (1998) (Fig. 1.132). Quartz is the most abundant mineral occurring throughout the vein. Adularia tends to occur at earlier stage. Smectite is the earliest mineral. Electrum tends to occur in early and middle stages. Naumannite occurs after the early-stage
"•,•quenceveil,of
I
t
a
miner~,.,
b
c
d
e
II f
g
h
111
i
Quartz
m
m
Adularia -i
~0
Smectite Electrum &"
.k"
:6_'
e6_
Naumannite
Ag-Sb-S "E ©
minerals
Pyrite Chalcopyrite Sphalerite m
abundant ~
rich
~
common-
poor - - f e w
Figure 1.132. Paragenetic sequence of the Hosen No. 5 vein (Takahashiet ai., 1998).
Miocene-Pliocene Hydrothermal Ore Deposits
187
electrum. Ag-Sb-S minerals are found both in early- and late-stages. Abundant pyrite occurs in early-stage. Small amounts of chalcopyrite are widely distributed. Sphalerite is found in late-stage. Stibnite and realgar are found in the latest stage. The sequence of mineralization from early to late is consistent with that of the other deposits (Asahi, Ogane, Takadama) sulfide stage (pyrite, marcasite, chalcopyrite, sphalerite, galena, electrum, naumannite, polybasitc, quartz, adularia), sulfosalt stage (pyrite, chalcopyrite, electrum, tetrahedrite, pyrargyrite, miargyrite) adularia, quartz (calcite), and barren stage (adularia, quartz, pyrite, realgar, stibnite) (Fig. 1.133). Adularia/quartz ratio decreases with the stage of mineralization and correlates with Au content of ore (Shikazono and Nagayama, 1993).
1.4.6.4. Geochemicalfeatures Oxygen and hydrogen isotopes. Oxygen and hydrogen isotopic studies on the altered rocks and gangue minerals were reported by Shikazono and Nagayama (1993), Naito et al. (1993), Matsuhisa and Aoki (1994), and Imai et al. (1998). Shikazono and Nagayama (1993) found that 8180 of adularia and quartz and adularia/quartz ratio decrease from early- to late-stage of the Hishikari mineralization, gold contents of ore samples in a vein positively correlate with 3180 and adularia/quartz ratio, as well as, the Pb, Cu, Ag and Se contents of vein samples, and Hg and As contents of the same samples increase towards the late-stage of mineralization (Fig. 1.104). Using fluid inclusion data and 3180 of adularia and quartz, it is inferred that 8180 of ore fluids decreases with the stage of mineralization (Shikazono and Nagayama, 1993). The early-stage ore fluids were of highly exchanged meteoric waters or magmatic water origin, while late-stage ore fluids were dominated by meteoric water (Shikazono and Nagayama, 1993). 8180 and 3D of ore fluids were influenced by clay minerals in the Shimanto Supergroup sedimentary rocks (Imai et al., 1998). The 813C value of H2CO3 and the 8~80 value of the hydrothermal solution using the data on vein calcite coexisting with electrum and assuming isotopic equilibrium with calcite and the temperature obtained by fluid inclusion microthermometry range from -14.4%o to -9.1%o, and from -6.2%o to +5.4%o, respectively, suggesting that the hydrothermal solutions isotopically equilibrated with the sedimentary basement rocks were responsible for the gold mineralization associated with calcite (Imai and Uto, 2001). Naito et al. (1993) showed that oxygen isotopic composition (81SO) of altered volcanic rocks in the Hishikari mine area varies systematically; +5.9%o to +15.9%o (zone I), +7.1%o to +12.4 %0 (zone II), +2.8 to +11.7 %o (zone III), and +2.1%o to +8.2 %o (zone IV) (Fig. 1.134). They calculated the change in 8280 values of hydrothermally altered volcanic rocks as a function of water to rock ratio by weight and temperature, assuming that oxygen isotopic equilibrium is attained in a closed system, and demonstrated that the increase in 8180 values of altered andesitic rocks from the veins towards peripheral zones can be interpreted as a decrease in temperature from the vein system (Fig. 1.135). In their calculations, the effect of mixing of hydrothermal solution with groundwater was not considered.
Chapter 1
! 88
Later
Earlier
Sulphide stage Sulphosalt stage Adularta Quartz Calcite Pyrite Marcasite Sphalerlte Chalcopyrlte Galena Gold Argentita Polybasite Tetrahedrlte Pyrargyrlte Mlargyrlte Realgar
mare
Barren stage
i
I
iwua=
I lib i
milm i
-? n
.,,?
Paragenetic sequence of the Asahi deposit Olta pref. Matsukuma (1951)
II Quartz Adularla Pyrite Chalcopyrlte Sphalerlte Galena Ag-mlnerele Gold Rhodochroalte Barite
~
III
-
~
--
-.-,
Psregenetic sequence of the Nishltanl vein, Nlahltanl vein group, the Ogane mine, Hokkaldo. Aklba (1957)
A
B
C
D
Quartz Adularla A:Big adularla crystals with fine quartz B:Aggregatee of adularla and quartz C:Masslve quartz with adulsrla l D:Drusy quartz General feature of paragenetlc sequence of gangue minerals, Tekatama mine, Fukushlma pref, modified from Yagyu (1954) Figure 1.133. "The normal order" in other epithermal gold deposits in Japan (Nagayama, 1993a).
Miocene-Pliocene Hydrothermal Ore Deposits
189
I 8.9+-1.4%-,n=50)
10
L.A. (least altered)
III IIII IIIIII
iI
Zone 110.9:1:2.7%.,n:~,l Cr-Sm 5 Cr-Ka(Ko)
6
llll I
i 1 i i I-I-I I11
I II
II
IK~KaI
5
lo
15
t I[ Zone (9.0±1.8%o, n=ll) Oz-Srn 5 Qz-Ko(Ko)
6
g
I"l R - F ~
Zone ( 6.9-+ 1,8% , n=32} 5 t Int.ClaYmirerolc
6
/ CM-Ser l |/
. o
I-13
I I III
5
I I I
Vl~ lo
VI
1~
I , r-I I-I
lo
1is
[7 Illllll [-'1 i i Ii Ii Ii Ii Ii 'i 'i I I I rl s
[-1
lb
1~
a~8o (%0) Figure I.i34. Histograms of 3[80 values of the volcanic rocks for five distinct alteration zones (Naito et aI., 1993). Shikazono et al. (2002) interpreted the ~180 zonation based on hydrothermal solution-groundwater mixing model. For the calculations the following assumptions were given: (1) Hydrothermal system at discharge zone is composed of five reservoirs such as ore deposit/zone IV boundary, zone III/II boundary, zone I I / I boundary, zone I/fresh country rocks boundary and temperature of each reservoir is 250°C, 220°C, 150°C, 100°C, and 25 °C, respectively.
Chapter 1
190
+20.C
250C-
+15. O
~
-
65°C
O 0 +10.0 ,.,f
100°(:;_ 113°C_
0
t43
1500C +5.0
200°C 250°(3 300°(3 I
i
i
I
i
1
I
1
2
3
4
5
6
7
9
Water/Rockratioin Weight Figure 1.I35. Change in 5180 values of hydrothermally altered rocks as a function of water to rock ratio with several different temperatures. The initial 3180 values of rock and water were taken as +8.5%o and -5.0%o, respectively (Naito et al., 1993).
(2) Initial 3180 of hydrothermal solution and groundwater are 0%0 and -7%0, respectively. (3) Oxygen isotopic equilibrium between mixed fluid and alteration minerals is attained. (4) Minerals in oxygen isotopic equilibrium with mixed fluid are feldspar for ore deposit/zone IV, and zone III/II, montmorillonite, and kaolinite for zone II/I, and montmorillonite for zone I/fresh rocks. Oxygen isotopic fractionation factors used for the calculation were taken from Taylor (1997). Initial 8180 value of hydrothermal solution (0%0) was estimated from 8180 values of K-feldspar and quartz in the veins and homogenization temperatures (Shikazono and Nagayama, 1993), and that of groundwater (-7%0) was estimated from meteoric water value of the south Kyushu district (-7%0) (Matsubaya et al., 1975). The mixing ratio of hydrothermal solution and groundwater was calculated based on the temperature of each reservoir. Using the estimated 81SO of fluid and oxygen isotopic fractionation between water and mineral, 3180 of altered rocks were estimated. The calculated result, together with the average 8180 of hydrothermal alteration zone by Naito et al. (1993), is shown in Fig. 1.136. This shows a fairly good agreement
Miocene-Pliocene HydrothermaI Ore Deposits
191
16
X 14
- • - Analytical value
12
~
--X-- Calculated value
.~ ~o
.....
o
£
• x. , -~~ . ~
,;o
,;o
=;o
=;o
aoo
Temp.(*(;)
Figure 1.136. 3180 change due to the mixing of hydrothermal solution and groundwater (Shikazono et al., 2002). between the model calculation and analytical data on 8180 except for the data on the most peripheral zone, suggesting the mixing of hydrothermal solution and groundwater is important for causing of 8180 variation in hydrothermally altered andesite.
~34S of sulfides. 834S values of sulfides in the veins are -1%o to +2%0 (Shikazono, unpublished). This is very close to average 834S value of Japanese epithermal A u - A g vein-type deposits which is ca. +2%o. In contrast to the 834S of hydrothermat solution for the vein, that of pyrite in hydrothermally altered rocks (Shimanto Shale) varies very widely, ranging from -5%o to +15%o. Based on the microscopic observation, pyrite with low 834S values less than 0%o is usually framboidal in form, suggesting that low 834S was caused by bacterial reduction of seawater sulfate. There are two possible interpretations of high 834S values (+10%o to 4-15%o). One is the reduction of seawater sulfate in a relatively closed system. The other one is a contribution of volcanic SO2 gas. As noted already, volcanic SO2 gas interacts with H 2 0 to form H2SO4 and H2S. 834S value of SO 2- formed by this reaction is generally high although the 834S value depends on the sulfur isotopic fractionation between H2SO4 and H2S and temperature. It is worth noting that in both cases a considerable degree of sulfate reduction took place. This may be due to the small water/rock ratio in the altered rock system compared to open crack systems through which large amounts of hydrothermal solution migrate. Bulk compositions of altered andesite. Bulk compositions of altered andesite were obtained by XRF analysis (X-ray fluorescence analysis) (Shikazono et al., 2002). Figures 1.137 and 1.138 show the change in elemental contents (SiO2, K20, CaO, MgO) of altered andesitic rocks away from the quartz vein system. Figure 1.139 shows the relationships between (CaO + Na20) content and K 2 0 content. These data indicate the following features of compositional variation of altered rocks.
I'o
90
B
A
8o
°l
C
l • 91i !
8t-
.
70
• 0 •
60 i ~5o
A
•
• •
.,,
••
~_ _
7t-
•
~8 •
"
•
5F
.0_40
i •
,, • •
.
4!
1 30
4
3 i-
•
20
•
2
•
•
2["
•
•
i
A
•
•
10
ti 0
240 4()0 660 D i s t a n c e l r o m the vein (m)
800
00
200 400 6()0 800 D i s t a n c e from the vein (m)
0 oi
260
460
600
~00
D i s t a n c e f r o m the vein (m)
Figure 1.137, The variations of major element contents in andesite (drilling core, 8-MAHAK-4 and underground samples) away from the vein system (Shikazaoo et al., 2002), Diamond: Hishikari Lower Andesitic tuff (underground samples): square: Hishikari Lower Andesitic tuff (drilling core samples); triangle: Hishikari Lower Andesitic lava (drilling core samples); ×: relatively fresh Hishikari Lower Andesitic lava (drilling core samples). (A) SiO2 content variation. (B) K20 content variation. (C) CaO content variation.
193
Miocene-Pliocene Hydrothermal Ore Deposits 4
3.5 3 o~ 2.5
*
|
2*
4'
•
0
if2 1.5 1 0.5 0
0
Distance from the vein (m)
Figure 1.138. MgO content variation for underground samples. Diamond: Yusen No. 7 vein; square: Seisen No. 8 vein.
14
12
10
1( "x ×
•
~8 o ll•
Z +
Abbreviations: A zone I, × zone II, , zone III, • zone tV, + z o n e IV ( n e a r the quartz vein), • fresh andesite (Hishikari Lower Andesite)
O6 o
AA
X &
2
• • • ÷
÷
•
+%
÷ #
÷
04.+~
++'~÷
÷
o
0
5 K20 (wt%)lO
15
Figure 1.139. The relationship between K20 content and (Na20 + C a • ) content in andesite (Shikazono et al., 2002).
Chapter I
194
(1) SiO2 content of andesite tends to be high near the veins. High SiO2 content may be due to the presence of quartz veinlets (Fig. 1.137A). (2) K20 and MgO contents of altered andesite decreases away from the vein (Fig. 1.137B and Fig. 1.138). (3) Analytical data on andesite are plotted on (CaO 4- N a 2 0 ) - K 2 0 diagram (Fig. 1.139). This shows that (CaO 4- Na20) content inversely correlates with K20 content. These results are consistent with XRD (X-ray diffraction) results. The amounts of K-feldspar, K-mica and chlorite are higher in the altered rocks closer to the veins and Ca-zeolites and smectite decrease in amounts towards periphery of the alteration zones. Although data are scattered and mostly no systematic correlation of Na20, CaO and MgO contents with 8t80 exist, K 2 0 content and 8~80 seem negatively correlated; K20 content is low and high for zone I and IV and 8180 is high and low for zone ! and IV, respectively. This correlation suggests that mineralogical alteration zoning correlates to 8180 variation that has been cited by Naito et al. (1993). The variations in K20, Na20 and CaO contents could be interpreted by thermodynamic consideration. The dependence of concentration of K +, Na +, Ca 2+ and H4SiO4 in equilibrium with common alteration minerals (K-feldspar, Na-feldspar, quartz) on temperature is shown in Fig. 1.140 (Shikazono, 1988b). This figure demonstrates that (1) chemical compositions of hydrothermal solution depend on alteration minerals, temperature and chloride concentration, and K + and H 4 S i O 4 concentrations increase and Ca 2+ concentration decrease with increasing of temperature. In this case, it is considered that potassic alteration adjacent to the gold-quartz veins occurs when hydrothermal solution initially in
E ..go
B
Na+
H
"
Ca2+
.A N
-1
-2
-3
1;0
200
2;0
300
Temp. (°C)
Figure 1.140. The dependence of concentration of K +, Na +, Ca 2+ and H4SiO4 in equilibrium with common alteration minerals (K-feldspar, Na-feldspar, quartz) with temperature (Shikazono, 1988b). Thermochemical data used for the calculations are from Helgeson (1969). Calculation method is given in Shikazono (1978a). Chloride concentration in hydrothermal solution is assumed to be 1 mol/kg H20. A-B: Na + concentration in solution in equilibrium with low albite and adularia. C-D: K + concentration in solution in equilibrium with low albite and adularia. E-F: H4SiO4 concentration in solution in equilibrium with quartz. G-H: Ca 2+ concentration in solution in equilibrium with low albite and anorthite.
Miocene-Pliocene Hydrothermal Ore Deposits
195
Total mKCJmHCI 600 0 500 o
400
Q. 300 E 200
Andalusite~ ~. ..... ~
K-feldspar
Pyrophylli~__ ~ ~
H~A
Kaolinite
100
Acidic GW 0
2 4 Total mKcJmHcI
GW I 6
Figure 1.141. Temperalure-aK+/aH+(a: activity) trend due to the mixing of hydrothermal solution and groundwater accompanied by hydrothermal alteration (Hemley and Jones, 1964). HS: hydrothermal solution, GW: groundwater, A: hydrothermaIsolution-groundwater mixing line, B: hydrothermal solution acidic groundwater mixing line (Shikazonoet al., 2002). equilibrium with propylitic alteration minerals (albite, K-feldspar, quartz) ascends rapidly and interacts with country rocks at lower temperature ( I - J - K in Fig. 1.140). It is also considered that K + is added from hydrothermal solution to the rocks as K-mineral such as K-feldspar and K-mica at the site of ore deposition, accompanied by the destruction of plagioclase in the country rocks and liberation of Ca 2+ and Na + to the fluid. It is expected that SiO2 content of the country rocks increases with proceeding of the alteration because solubility of SiO2 decreases with a decrease in temperature ( O - P - Q in Fig. 1.140). Figure 1.141 shows that the hydrothermal alteration zoning from K-feldspar through K-mica to kaolinite from the vein towards marginal part of alteration zones can be explained by a combined change of temperature and m~+/mH+ (m: molal concentration) ratio. Decreases in temperature and mK+/mH+ ratio of hydrothermal solution cause the above alteration zoning. This trend suggests that ascending hydrothermal solution mixed with acidic (low pH) and low-temperature descending groundwater (B in Fig. 1.141). Acidic groundwater may be formed by the interaction of groundwater with acidic alteration zone at shallow part, input of volcanic SO2 gas into groundwater, condensation of steam, or oxidation of H2S. If K + is added to the rock accompanied by the destruction of feldspar, the following reactions occur. CaO (Ca-feldspar) + 2 K + ~ K 2 0 (K-feldspar) + Ca 2+
(1-56)
Na20 (Na-feldspar) ÷ 2 K + --+ K 2 0 (K-feldspar) + 2 Na +
(1-57)
It is expected from (1-56) and (1-57) that CaO and Na20 contents of altered rocks inversely correlate with K20 content of altered rocks with a negative slope of - 1 . Analytical results show the negative correlation on (CaO + N a 2 0 ) - K 2 0 diagram
Chapter 1
196
(Fig. 1.139). However, the low CaO 4- Na20 content and K20 content data on Fig. 1.139 cannot be explained only by these reactions. The relatively low contents data could be attributed to the dissolution of silicates by acidic solution which was generated by the disproportionation reaction, 4SO2 (volcanic gas) + 4 H 2 0 ~ 6H + + 3 SO 2- + HzS
(1-58)
The following dissolution of silicates such as feldspar occur by the interaction of H + with feldspar, CaO (Ca-feldspar) + 2H + ~ Ca 2+ 4- H20
(1-59)
Na20 (Na-feldspar) 4- 2H + ~ 2Na + 4- H20
(1-60)
K20 (K-feldspar) 4- 2H +
(1-61)
--+ 2K + 4- H20
Therefore, it is thought that the mixing of acidic solution with hydrothermal solution occurred and andesite near the gold-quartz veins suffered superimposed potassic and advanced argillic alterations.
1.4.6.5. Interpretation of Si02 mineral zoning in terms of kinetics-fluidflow mixing model
Shikazono et al. (2002) considered the depositional mechanism of quartz and cristobalite and the change in silica concentration of fluid migrating through the altered rocks in the Hishikari mine district based on kinetics-fluid flow mixing model. Their discussion is summarized below. Hydrothermal solution containing appreciable amounts of dissolved silica migrates through andesitic volcanic rocks, accompanying SiO2 precipitation. Figure 1.142 shows the dependence of solubility of SiO2 minerals (quartz, cristobalite) on temperature. As described already, cristobalite occurs in peripheral and shallower part of hydrothermal alteration zone. Quartz is present in zones occurring in deeper and closer to the gold-quartz veins. Such zoning from quartz to cristobalite is also common in main active geothermal systems (Hayashi, 1973; Takeno et al., 2000). Precipitations of quartz and cristobalite occur due to a decrease in temperature that is caused by heat conduction, and fluid mixing. The changes in the concentration of dissolved silica during these processes are shown in Fig. 1.142. Decrease in temperature due to heat conduction alone cannot explain the distributions of quartz and cristobalite. The temperature at which the heat conduction trend crosses the cristobalite saturation curve is ca. 200°C, which is higher than the 100°C corresponding to the cristobalite/quartz boundary in active geothermal system (Hayashi, 1973; Takeno et al., 2000). The curve for the mixing of hydrothermal solution and groundwater always lies below the cristobalite saturation curve if the hydrothermal solution is in equilibrium with quartz. Therefore, the heat conduction and mixing of fluids in equilibrium with quartz are considered to be not main causes for the precipitation of quartz and cristobalite. Therefore, in order to know the change in dissolved silica concentration and temperature during the precipitation of quartz and cristobalite and mixing of fluids, the following equation could be used:
Miocene-Pliocene Hydrothermal Ore Deposits 800
197
800i
i (B)
(A) 700
700!
[]
E~ 0
o
600
6OO
[]
[3
~
0 H.S.--
500
400
[] H.ST"-
500
4O0
g o5
8 O300
300'
200 i
200
~
I 100
~.~.
/
._~ ~ 2
~
100 150 Temp (°C)
200
250
0.1
A / M ~ 1.00
l°° i ~ D D - ^ A
50
~/~
[3
o
%ee~%5
AA-
~
550
200
250
Temp (°C)
Figure 1.142. The computed result of the relationship between dissolved silica (H4SiO4) concentration of mixed fluid and temperature based on four reservoirs model (Shikazono et aI., 2002). Open triangle: solubility curve for quartz, Open square: solubility curve for c~-cristabalite,Solid triangle: Hishikari Lower Andesite lava (drilling core), Cross: Relatively fresh Hishikari Lower Andesite lava (drilling core). H.S.: hydrothermal solution; G.W.:ground water.
dC/dt = k(A/M)(Co - C) ÷ ( q l C 1 / V ) + (q2C2/V) - (ql + q2) C~ V
(1-62)
where C = concentration o f dissolved silica of output fluid ( m o l / k g H 2 0 ) , t = time (s), k = rate constant (mol m - 2 s 1), Co = concentration of dissolved silica saturated with respect to quartz ( m o l / k g H 2 0 ) , C1 = concentration of H4SiO4 o f input hydrothermal solution ( m o l / k g H 2 0 ) , C2 = concentration of H4SiO4 of input groundwater ( m o l / k g H 2 0 ) , ql = volume flow rate of hydrothermal solution (m 3/s), q2 = volume flow rate of groundwater (m3/s), M = mass of aqueous solution in a system (kg), V = volume of aqueous solution in a system (m3), and A = surface area of rocks which contacts with aqueous solution (m2). If steady state is attained, we obtain,
C = {k(a/M)Co + (ql C1 + q2C2/V } / { k ( a / M ) + (ql + q2)/V }
(1-63)
As shown in Fig. 1.143, the system is divided into four reservoirs. Each reservoir corresponding to alteration zone IV, III, II and I is assumed to be homogeneous with respect to temperature and concentrations of dissolved silica in aqueous solution. The temperature of the initial hydrothermal solution is assumed to be 250°C from homogenization temperature of fluid inclusions in vein quartz (Shikazono and Nagayama, 1993). Temperature o f each reservoir was estimated from the assemblage of hydrothermal alteration minerals and temperature of alteration zone in active geothermal system (e.g., Hayashi, 1973; Takeno et al., 2000).
Chapter 1
198 G°W.
H.S. ~ T ~250~
T ----220
T ~150
[ T ~-100
T =25
Figure 1.143. Four reservoirs model. H.S.: hydrothermal solution, G.W.: groundwater (Shikazono et at., 2002).
Volume flow rates of incoming hydrothermal solution and groundwater to each reservoir were estimated from the temperature of reservoirs and initial hydrothermal solution (250°C) and groundwater (25°C) Volume of solution (V) in a reservoir is expressed as, V = Vbox~b/100
(1-64)
where Vbox = volume of reservoir as a rock, and q~ = porosity. Volume flow rate (q) is expressed as, q = A'v(a/lO0
(1-65)
where A I = cross-section area of box (m 2) and v = velocity of fluids. It is assumed that the following simple relation is approximately established in a case of two fluids (fluid 1 and fluid 2) mixing. ql 7"i ÷qzT2 = (qI + q z ) T
(1-66)
where T = absolute temperature. The mixing ratio for fluid (R) is approximated as, R = q2/(ql + q2)
(1-67)
The calculations based on four reservoir models were made using equations (162)-(1-67) and precipitation rate constant (k) for SiO2 minerals by Rimstidt and Barnes (1980). Volume of each reservoir was calculated from the volume of each alteration zone and assuming porosity of alteration zone to be in a range of 3-1%. Dissolved silica concentration of groundwater was assumed to be 60 mg/kgH20 that is average dissolved silica concentration of groundwater in andesitic volcanic region in Japan (Shikazono, unpublished). Assuming the ranges of A / M , flow rate of mixed fluid, porosity and giving the precipitation rate of SiO2 minerals (Rimstidt and Barnes, 1980), the relationship between dissolved silica concentration of mixed fluid and temperature was obtained (Fig. 1.142). It was found that the porosity does not change the results of calculations. Figure 1.144 shows the results of calculation based on multireservoirs (40 reservoirs) model in which each reservoir corresponding to each alteration zone is divided into
Miocene-Pliocene Hydrothermal Ore Deposits
199
800
800
700
700 :
[] [2
600
E3 E3
600 []
[] C] H.S.-- ,
5O0
~ 400
[] H.S.~
400 [3 °
8
G
500
oo
G
b5 3OO
300 A/M~0.0
V = l O -3 ~
__10~.2
200
100200 G
~
I
°~'a
-4¢dM~1.00
1O0
50
100
150
Temp (°C)
200
250
50
100 150 Temp (°C)
200
250
Figure t. 144. Computedresults for multireservoirs (40) model. Open triangle: solubility curve for quartz, Open square: solubility curve for c~-cristabalite, Cross: no precipitation, Open circle: computedresult (Shikazono et al., 2002). ten reservoirs and temperature of each reservoir was given. The results of calculation indicate that the flow rate, 10.4.2 m/s, is the best estimate for A/M = 0.1 that is estimated from the width of quartz veinlets in altered andesite, and this rate is in agreement with the flow rate of geothermal water in active geothermal system ( 1 0 - 6 - t 0 .4 m/s) (Fujimoto, 1987). In this model, the area considered is (2 km x 0.5 km) and flow rate is 10 -4.2 m/s. When porosity is 2%, total mass flow rate is estimated as 10 42 m x 106 m 2 × 2/100 = 1.4 × 106 g/s. This is very similar to that in the Wairakei geothermal area in New Zealand which is 1.3 x 106 g/s (Elder, 1966).
1.4.6.6. Goldprecipitation due to mixing of fluids in epithermal system The mineralogy of hydrothermal alteration zoning, bulk compositional variation of altered rocks, thermodynamic consideration, mass transfer, and oxygen isotope computations, and the sulfur isotope study mentioned above all suggest that the mixing of two fluids (hydrothermal solution and acid groundwater) is the main cause for the geochemical and mineralogical variations in the Hishikari gold mine district. A large number of studies on the depositional mechanism of gold in epithermal system have been carried out (e.g., Shikazono, 1974a, 1986; Shikazono et al., 1990, 1993; Drummond, 1981; Reed and Spycher, 1985; Spycher and Reed, 1989). Depositional mechanism of gold in the Hishikari deposit has been discussed by several investigators (Izawa, 1988; Izawa et al., 1990; Shikazono and Nagayama, 1993; Nagayama, 1993b; Hayashi et al., 2000a,b). For example, Izawa (1988) and Izawa et al. (1990) thought that
200
Chapter 1
the mixing of ore fluids with groundwater, boiling and oxidation of ore fluids due to the interaction of ore fluids with oxidized hematite-rich paleosol are the main causes for the deposition and enrichment of gold in the veins. Hayashi et al. (2000a,b) suggested from a drastic change in 3180 of quartz at the precipitation sequence that electrum precipitated due to the mixing of fluids. The previous studies clearly demonstrated that Au thio complex is dominant among dissolved Au species in ore fluids responsible for Japanese epithermal A u - A g vein-type deposits (e.g., Shikazono, 1974a), considering the estimated f Q - p H range of Japanese epithermal gold deposits and the gold solubility due to thio complex (Seward, 1973, 1981). Thus, accepting the assumption that the Au thio complex is dominant among dissolved Au species in the ore fluids responsible for the Hishikari deposit, the precipitation reaction of gold in electrum is expressed as, Au(HS) 2 + H + + 1/2H2 ~ Au + 2H2S
(1-68)
This reaction suggests that a decrease in H2S concentration, and increases in H + concentration and fH2 (H2 fugacity) and temperature variations are important causes for the deposition of gold in electrum. It is also likely that the deposition of gold in electrum occurs by the following oxidation reaction. Au(HS) 2 + 1 5 / 2 0 2 + 1 / 2 H 2 0 --+ Au+2SO42 + 3 H +
(1-69)
This reaction suggests that oxidation of fluids is important as a depositional mechanism. However, this oxidation reaction seems difficult to explain the gold deposition from the following reasons. (1) The rate of oxidation of H2S to S O ] - is slow at a site of gold deposition. (2) O2 concentration of groundwater is very low. (3) HzS concentration in epithermal ore fluids seems higher than SO 2-. It is also possible that the following reaction is important for the deposition of gold. Au(HS)2 + 8 H20 ~ Au + 2 SO,]- + 3 H + + 15/2 H2
(1-70)
This reaction proceeds due to the degassing of H2, decreasing of H + and probably decreasing of temperature. We cannot evaluate this reaction as an important cause for gold deposition because no study on this reaction has been done. The importance of reaction (1-68) has been already cited by Reed and Spycher (1985) and Shikazono and Nagayama (1993). Shikazono and Nagayama (1993) favored the two fluids mixing as a cause for the gold deposition by the reaction (1-68) from the presence of acidic alteration zone overlying the Hishikari gold-quartz veins. For example, alunite and silicified rocks occur in higher elevations in the north-eastern part of the Hishikari district (Izawa et al., 1990). Shikazono (1985a) suggested based on the studies of wall rock alterations associated with Japanese epithermal-type Au-Ag deposits that acid sulfate fluids coexist with near-neutral chloride-rich fluids at relatively shallow zone from the surface (less than 1 kin) at the time of epithermal-type A u - A g ore formation. Reed and Spycher (1985) made a computation on the gold deposition from the mixed fluids. Their results support that gold deposition occurred in the Hishikari veins
Miocene-Pliocene Hydrothermal Ore Deposits
201
due to the mixing of two fluids (hydrothermal solution and acidic sulfate groundwater). However, it is uncertain that the alteration minerals (K-feldspar, K-mica and kaolinite) formed in the same stage of hydrothermal alteration. It may be possible that descending acid low-temperature solution (groundwater) formed kaolinite and cristobalite at different stage of gold-quartz mineralization associated with K-feldspar precipitation (Hedenquist et al., 1996). If such kind of mixing occurs, pH decreases, H2S concentration decreases, fH2 decreases and temperature decreases. These changes except the decrease of H2 proceed reaction (1-68). Another important depositional mechanism for gold in electrum is boiling of ore fluids, as inferred by many investigators (e.g., Drummond, 1981). Boiling of ore fluid causes an increase in pH. If dominant Ag species is AgCIf, this pH change as well as decreasing temperature, H2 degassing and dilution (decreasing of C1 concentration) causes the deposition of Ag in electrum according to the following reaction. AgC12 + H20 --+ Ag + 2 C1- + H + + 1/2 H2
( 1-71)
Galena, sphalerite and chalcopyrite precipitate also due to increase of pH due to the breakdown of base metal chloro complexes accompanied by the boiling by the reactions such as, PbC12 + H2S --+ PbS + 2C1- + 2H +
(1-72)
ZnC12 + H2S --+ ZnS + 2C1- ÷ 2H +
(1-73)
CuC12 + 2H2S + FeC12 --+ CuFeS2 + 4C1 + 4H +
(1-74)
However, in order to clarify the depositional mechanism of electrum and sulfides, more detailed description of alteration minerals, 3180, 3D data and the salinity (C1concentration)-enthalpy relationship are clearly required, and the two fluids mixing model has to be evaluated based on these data. It is difficult to evaluate that the boiling is an important mechanism for depositions of sulfides and electrum because of scarce data on fluid inclusions (salinity, temperature, enthalpy, 3180 and ~D). However, the above consideration and previous studies on epithermal Au-Ag deposits stressed an importance of boiling for gold depositions (Seward, 1991; Hayashi et al., 2000b). Therefore, it cannot be ruled out that the boiling and H2S loss are main causes for the depositions of electrum and sulfides (spbalerite, galena, and chalcopyrite) in the Hishikari deposit. However, preliminary study on the fluid inclusions indicates that boiling zone and ore zone containing high gold content are different (Etoh et al., 2001). Probably, the mixing of boiled fluid with acid groundwater caused efficient deposition of electrum in the Hishikari hydrothermal system.
1.5. Evolution of tectonics and hydrothermal system associated with epithermal and Kuroko mineralizations Numerous studies on the geologic and tectonic evolution in and around the Japanese Islands from Miocene to present have been carried out (e.g., Kitamura, 1959;
Chapter 1
202
Ozawa, t963; Sugimura and Uyeda, 1973). For example, Sugimura and Uyeda (1973) summarized volcanic rocks, degree of deformation of sediments, structural trends of sediments deposited, and amount of uplift and subsidence since Miocene. However, in contrast to these geologic and tectonic studies, very few studies on the relationship between tectonics and hydrothermal system in Neogene age have been carried out. Therefore, these studies are briefly summarized and then the relationship between geologic and tectonic evolution and evolution of hydrothermal system associated with the mineralizations (Kuroko deposits, epithermal veins) are considered below. 1.5.1. P a l e o g e o g r a p h y and stress field
Geologic environments and tectonic evolution of Green tuff region in Northeast Honshu have been studied by many workers (e.g., Kitamura, 1959). Paleogeography of the Japanese Islands and surrounding areas during a period of middle Miocene age (Kuroko-stage) and late Miocene-Pliocene age (vein-type stage) have been reconstructed in Fig. 1.145, mainly based on the paleontological studies (e.g., Ikebe, 1973, 1978; Chinzei, 1991). According to these studies, Green tuff region of middle Miocene age has been submarine environment in most areas (Fig. 1.145). Rapid subsidence occurred at middle Miocene age, which was contemporaneous with the age of Kuroko mineralization (Yamaji, 1990). From about 5 4-2 Ma uplift of Green tuff region in Inner Honshu province, Northeast Japan took place and this region became subaerial (Sugi et al., 1983; Otsuki, 1989, 1990).
(a)
(b)
o
K T 1 o*
,,¢(~
; @ushO
,.
L.V
0
200km
Land area -, -,~KKfi;hu Y £;k Coastal line *,v ~ ' ~,' (present-day) "~""
~3 Land area I
0
I
O
200km
Coastal line (present-day)
Figure 1.145. (a) Middle Miocene (Kuroko-stage)(solid circle: Kurokodeposits, open circle: Au-Ag vein-type deposits). (b) Late Miocene-Pliocene (vein-stage)(open circle: Au Ag vein-typedeposits) (Shikazono, 1987b).
Miocene-Pliocene Hydrothermal Ore Deposits
203
It is worth noting that the age of uplift is nearly coincident with the age of epithermal vein-type deposits in Northeast Japan including the Chitose, Yatani, Takadama, and Nebazawa Au-Ag vein-type deposits and Hosokura Pb-Zn vein-type deposits. However, the ages of some large base-metal vein-type deposits in Northeast Honshu (Ani, Osarizawa and Nikko Cu deposits) are older than this uplift, suggesting that unconsolidated marine sediments containing interstitial seawater were distributed in the mine area at that time. 334S data on sulfides and sulfates and high salinity of fluid inclusions from these large base-metal vein-type deposits support the above argument. Unfortunately tectonic situations of the regions other than Northeast Honshu of Neogene age are not well understood. However, it seems evident that even in the regions other than Northeast Honshu epithermal Au-Ag vein-type deposits formed when the uplift started and the area of land expanded. In addition to the paleontologic data, the country rocks of epithermal Au-Ag mine districts also suggest that epithermal Au-Ag vein-type deposits have formed under the subaerial condition: welded tuff occasionally occurs in the mine area (e.g., Sado, Nebazawa, Northeast Hokkaido) and in general submarine sedimentary rocks and volcanic rocks are poor or absent in the Au-Ag mine districts (e.g., epithermal Au-Ag vein-type deposits in Kyushu). As already noted, most epithermal Au-Ag vein-type deposits are hosted by young (late Miocene-Pliocene) volcanic rocks and by sedimentary rocks, but dominant host and country rocks for base-metal vein-type deposits are submarine sedimentary and volcanic rocks. Submarine felsic tuft', tuff breccia, dacite lava, intrusive rocks and mudstone are dominant host and country rocks of Kuroko deposits. Detailed studies on the change of stress field during Miocene to present in Northeast Japan have been carried out based on (1) the measurements of direction of dike (Nakamura, 1977; Kobayashi, 1979; Takeuchi, 1980, 1981, 1987), and (2) measurements of direction and age of epithermal vein-type deposits (Horikoshi, 1975b; Otsuki, 1989). Otsuki (1989) recognized the Neogene tectonic stress fields of the Northeast Honshu Arc by analyzing the epithermal veins and showed that in middle to late Miocene, ENE-trending al and a2 coexisted and a3 had a NNW trend, and it was replaced by E - W compression at 7 Ma in Southwest Hokkaido and at 5 Ma in Northeast Honshu. This view is supported by the analysis of trend of dikes (Tsunakawa and Takeuchi, 1986; Takeuchi, 1987) which indicates that the stress field changed from extensional to compressional during 8-6 Ma in southern Northeast Honshu (Figs. 1.146 and 1.147). These analyses of stress field suggest that the dip of subduction of Pacific plate during Miocene might have been steeper than during Plio-Pleistocene considering the mode of subduction of Pacific plate in Northeast Honshu (Niitsuma, 1979; Niitsuma and Akiba, 1984). Usually active plate subduction with gentle dip causes uplift of land and expansion of land area (Uyeda and Kanamori, 1979). Such change from steep subduction to a gentle one caused the changes in site of hydrothermal activity from submarine area during Miocene (Kuroko mineralization) to subaerial area by Plio-Pleistocene to present (epithermal vein-type mineralization). It is also noteworthy that major igneous and hydrothermal activities in the Japanese areas seems likely to have taken place almost at same times as the stress changes. The stress changes occurred at about 22 Ma, 15 Ma, 12 Ma and 8 Ma in the southern part of
Chapter I
204
3)
2)
...........
;~...... ..................... L .....
~
3
Ma
22~18
16-12 N
(I) -- ~fl
i' ' .;! ~
p , 6~(.~i.;....,,!S T11
0I
200km 1 TRENDOF DIKES NORMALFAULT •'~ REVERSEFAULT STRIKESLIPFLT. ~.','~,~STRESSTRAJECT.
12-8
6-(0)
Figure 1.146. Stress trajectory maps of southern Northeast Honshu in the late Cenozoic period, after Tsunakawa and Takeuchi (1986) with a slight addition. C~H...... trajectory is drawn by smoothing the inferred stress orientations from the selected dike-swarms with K Ar dates. Selected major faults with age estimation are also shown for indicating types of stress fields. T: Extensional stress field, where O'v > O-Hm,x >> OHma., and normal or gravity faulting is preferable. P: CompressionaI. O-H,,,~,~>2> CrHm,,~> OV, reverse or thrust faulting (Takeuchi, 1987).
Northeast Japan and at about 15 Ma in the eastern part of Southwest Japan (Tsunakawa, 1986; Takeuchi, 1987) (Fig. I. 146). The changes in stress fields, and intensities of igneous and hydrothermal activities seem to correlate to oscillatory motion of the Pacific plate (Jackson's episodes) (Jackson et al., 1975; Jackson and Shaw, 1975) (Masuda, 1984). Masuda (1984) and Takeuchi (1987) pointed out that the oscillatory motion of Pacific plate during the least 42 Ma correlates with magmatism, the intensity of tectonism, the change of stress field and the history of sedimentary basin in arc-trench system (Fig. 1.147). The above arguments also suggest that the mineralizations in arc and back-arc systems relate to the oscillatory motion of the Pacific plate.
1.5.2. Volcanic activity The type of volcanic activity in and around the Japanese Islands changed throughout Tertiary. In early Tertiary subaerial andesitic activity was intense. For example, in
Miocene-Pliocene Hydrothermal Ore Deposits AGE
EO.
Ma BP
40
OLIGOCENE
EARLYMIOCENE MIDDLE MIOCENE LATEMIO. PLIOCENEiQUAT.
30
Hltl
20
I"
15
'I'"
I
I
cSua
10
5
tl /1"'
"
]
ci
~
205
I
1.0 . 5
I
II/ll II
c2
d
~ AVE,AoErH
NO
OF HAWAIIANCHAIN
.~ ,-,
90t
cou,to,~lO~,,i,~
Episodes
0
o [ I
Inner Stress Field
r Hmax ORIENTATION
T
n .
"
(~n(P)
r .
"
P transverse
~
i
]
Figure 1.147. Jackson's curve and arc stress reorientations. Apparent swing motion of Pacific Plate (Jackson et al., 1975) and regional stress orientation at the Northeast Honshu convergent margin are illustrated in order to show their synchronous relationship. Dashed line represents the average trend of the Hawaiian volcanic chain. Pacific plate moves along the direction with fluctuation in reference to Hawaii Hot Spot. Vertically shaded parts of the graph indicate the climax phases of "clockwise episodes". Lower part of the figure shows the phases and reversals in orientation of tectonic stress fields on the inner zone of Northeast Honshu Arc (Takeuchi, 1987). Northeast Honshu, subaerial andesitic volcanic activity was dominant at Daijima stage (mostly early Miocene). F r o m middle Miocene b i m o d a l basaltic-dacitic activity started with rapid subsidence in Northeast Japan. The production o f volcanic activity was probably greater than today. Basalt at middle Miocene age erupted at Northeast Japan was studied by Shuto (1989) and Tsuchiya (1988, 1989) who showed that basaltic m a g m a generated in deep mantle. Dudfis et al. (1983) showed that p r e - K u r o k o ore basalt in the Hokuroku district has Mg number ( M g O / ( M g O + FeO*)), where FeO* is total iron content expressed as FeO, in the range 0.85-0.67 4-0.01, suggesting that the basalt is a relatively primitive, unfractionated, and mantle-derived melt. From late Miocene to present, subaerial a r c - v o l c a n i c activity ( c a l c - a l k a l i rocks, andesite, tholeiitic and high alumina basalt) started associated with uplift of the Japanese Islands. This volcanic activity is different from that at middle Miocene age. The above-mentioned changes in paleogeography, volcanism, crustal movement (subsidence and uplift), and stress field clearly demonstrate that these features of back-arc volcanism in e a r l y - m i d d l e Miocene are quite different from those o f Island arc volcanism in late Miocene to present. According to Yoshida and Yamada (2001), the age of change in volcanism from back-arc type to Island arc type in Northeast Honshu was 12.7 M a and this age corresponds to the age of Kuroko formation.
206
Chapter I
1.5.3. Tectonic influence on temporal and spatial relationships in Kuroko and vein-type deposits in southern Hokkaido, Japan It is worth studying Kuroko and vein-type deposits occurring in one metallogenic province. Shikazono and Shimizu (1993) carried out integrated geological, geochemical and mineralogical studies on Kuroko and vein-type deposits in a southwest metallogenic province which is described below. Geology of the province is composed of Paleozoic basements, Tertiary altered submarine volcanic and sedimentary rocks (Green tuff) and Quaternary volcanic rocks. The basements are shale, tuff, limestone and chert of unknown ages. A simplified geologic map is shown in Fig. 1.148. Tertiary rocks are distributed widely. They are composed of alternations of sandstone, mudstone, andesitic and dacitic tuff, tuff breccia and lava. These rocks are intensively and extensively altered and are called as Green tuff. Tertiary volcanic rocks are variable in composition. Andesite, dacite and basalt are found. Quaternary volcanic rocks are dominantly andesite lava and are abundantly distributed in the northern part of the province (Fig. 1.148). The vein-type deposits can be divided into two based on the metals produced; precious (Au, Ag) and base metal (Pb, Zn, Ag, Mn, Cu, Fe) vein-types. There are two sub-types of the base metal vein-type deposits, the C u - P b - Z n sub-type and the P b - Z n - M n - A g sub-type. C u - P b - Z n veins occur in southern part of the province. Large P b - Z n - M n - A g veins and A u - A g veins are distributed in northeastern part. In the northeastern part, A u - A g vein-type deposits occur in marginal zones of the province, while the base metal-rich deposits ( P b - Z n - M n veins and Kuroko deposits) in central zone (Fig. 1.149). The marginal zone is characterized by exposure of Quaternary volcanic rocks and Plio-Pleistocene volcanic rocks in which Au-Ag veins occur, whereas the central zone is by thick submarine volcanic rocks (Fig. 1.150), in which base metal-rich deposits (base metal veins and Kuroko deposits) occur (Fig. 1.150). Tertiary volcanic rocks, Quaternary volcanic rocks and faults are distributed, trending generally from NW to SE. Some C u - P b - Z n veins in southern part are hosted by basement rocks. On the other hand, P b - Z n - M n - A g and Au-Ag veins occur in Tertiary and Quaternary volcanic rocks. Recently K - A t age dating of ore deposits, associated volcanic rocks and plutonic rocks, have been carried out. These data are summarized in Table 1.25. It is obvious that the Kuroko deposits have been formed at middle Miocene, being very similar to the ages of the Kuroko mineralization at Hokuroku district, northeast Honshu where large Kuroko deposits occur. On the other hand, the vein-type deposits were formed during Plio-Pleistocene ages. It is also worthwhile to note that the ages and distribution pattern of the vein-type mineralization in this province are very similar to those of andesite which overlies the vein-type deposits (Watanabe, 1990b; Sawai et al., 1992b). Although ages of volcanic rocks which host the Kuroko deposits have not been determined, it is obvious that the Kuroko mineralization took place at the time of formation of volcanic rocks (mainly dacite). Therefore, it is clear that felsic volcanic activities were related to the mineralizations of the Kuroko deposits, while andesitic volcanism to the vein-type deposits.
Miocene-Pliocene Hydrothermal Ore Deposits 140°E
I
207
141,E
N
Inakuraishi
l
Akaiwa vein
ocene ks
Imai-lshizoki ~km
Basemen[ Rocks
Figure 1.148. Simplified geologic map and distribution of the Kuroko-typeand vein-type deposits in south Hokkaido (Shikazono and Shimizu, 1993). Open circle: precious vein-type deposits. Solid circle: base metal vein-type deposits. Solid square: Kurokodeposits. Minor elements associated with the vein-type and Kuroko deposits are different. Characteristic minor elements concentrated to the ore deposits are Se, Te, Hg, As, Sb and Bi in the A u - A g vein-type deposits, Ag, Bi, As, Sb, Sn, W and Mo in the base metal vein-type deposits, and Au, Ag, Sb, As, Mo and Bi in the Kuroko deposits. This difference in minor elements is consistent with that found in the other epithermal vein-type deposits in Japan (Shikazono and Shimizu, 1992). Analytical results of sulfur isotope previously obtained are summarized in Fig. 1.151.
Chapter 1
208
°; \ o ,"o o Q "
\
\.
/
10oB,o"~,~ s A P P o R o
6 Hg
~
ell O"a~\-
Sn
AgSn"13
~%'Mo \
TeSe~~ °
/~~/
\.
k
',Bi \.
,4"
o°
o.o
• CuPbZnMnVeins~ O AuAgVeins • Kurokoor massivebaritedep.
•
/:%
•
0
18
'\
1119
J
\
\/
2"0 , . - " 1
\
Figure 1.149. Distribution of ore deposits in northeastern part of the province (modified after Yajima, 1979). 1: Suttsu, 2: Kutosan, 3: Toyoha, 4: Akaiwa, 5: Matsukura, 6: Meiji, 7: Todoroki, 8: Teine, 9: Kobetsuzawa, 10: Otoyo, 11: Inatoyo, 12: Toyohiro, 13: Jozankei, 14: Toyotomi, 15:Koryu, 16: Eniwa; 17: Chitose, 18: Shiraoi, 19: Morono, 20: Minami-Shiraoi (Shikazonoand Shimizu, 1993).
It is found that ~348 values for different types of ore deposits are different; 4-1.8%o to 4-5.1%o for the A u - A g vein-type deposits, -2.8%0 to +7.5%o for the base metal veintype deposits and +4%o to 4-5%o for Kuroko deposits. 334S of the base metal vein-type deposits is widely variable, but generally higher than that of the A u - A g vein-type and nearly similar to that of Kuroko deposits. It is worthwhile to note that average ~34S values for the base metal vein-type deposits (except for the deposits in basements), A u - A g metal vein-type and Kuroko deposits in the province are identical to those previously obtained for the other metallogenic provinces in Japan (Shikazono, 1987b). It is also found that 334S values of the ore deposits are related to the host rock types; 334S of ore deposits hosted by basement rocks are relatively low. For example, 334S of the Imaiishizaki and Sasayama is -2.1%o and -0.9%0, respectively. 334S values of relatively large ore deposits such as the Toyoha, Jokoku, and Ohe are high in a range of 4-4%~ to 4-8%0, compared with those of small ore deposits. It is likely that sulfide sulfur in ore fluids responsible for small ore deposits is influenced by the surrounding rocks having low 334S values.
Miocene-Pliocene Hydrothermal Ore Deposits
209 141 °
140 ° 43 °
i~.! ¥ Vv
Quaternary
Pliocene
Miocene
Pre-Tertiary
subaerial volcaniclastic deposits subaerial lava & volcaniclastic rocks coarse laminated sandstone, subaerial to submarine lava & volcaniclastic rocks (Setana E) massive mudstone, submarine lava & volcaniclastic rocks (Kuromatsunai F.) mudstone & shale, submarine lava & volcaniclastic rocks (Yakumo E) granitic rocks sandstone, submarine lavas & volcaniclastic rocks (Kunnui E) mudstone & non marine sediments (Yoshioka F.) subaerial lava & volcaniclastic rocks (Fukuyama F.) ~ basement . . . . . . . rocks
Figure 1.150. Simplified geologic map of northeastern part of the Hokkaido metatlogenic province (modified after Yamagishi, 1989) (Shikazono and Shimizu, 1993).
Available homogenization temperatures of fluid inclusion from the base metal vein-type, A u - A g vein-type, and Kuroko deposits are summarized in Fig. 1.152. Salinity (NaC1 equivalent concentration) of inclusion fluids is 1-6 wt%, 1-14.5 wt% and 0 - 3 wt% for Kuroko deposits, base metal vein-type deposits, and A u - A g vein-type deposits, respectively. These data clearly demonstrate that the salinity of inclusion fluids for the base metal-rich deposits (base metal vein-type deposits, Kuroko deposits) is higher than that of the A u - A g vein-type deposits, while homogenization temperatures of fluid inclusion for these three types of ore deposits do not show a wide
Chapter I
210
Precious veins (/3
_>, c-< o .13
E Z
-5 -4 -3 -2 -1 0
1 2 3 4
5 6
7 8 9 10
(~ 34 8 ( % 0 )
B a s e metal veins 03 0'1 t~ c.< 0 ..Q
E Z
, l;>4Xb
t
-5-4-3-2-10
1 234
5 678910
1~ 34 8 ( % 0 )
Kuroko >09, c--
< "5
..Q
E Z
,
1
L
i
-5-4-3-2-10
~
i
1 2 3 4
56
7 8
910
1~ 34 8 ( % o )
Figure 1.151. Sulfur isotopic compositions of sulfides in the vein-type and Kuroko deposits. Solid box represents sulfur isotopic data from the ore deposits occurring in basement rocks (Shikazono and Shimizu, 1993).
Miocene-Pliocene Hydrothermal Ore Deposits
211
TABLE 1.25 K-Ar age data on the vein-type and Kuroko-type deposits in southern Hokkaido (Shikazono and Shimizu, 1993) Name of mine Vein-type deposits Chitose Koryu Toyoha Ohe lnakuraishi Todoroki Teine Ofukeshi Yakumo Jokoku Shizukari
K-Ar age (Ma) 4.7 3.6 3.5, 3.4, 3.3 1.1 2.2 3.3 3.3, 3.4 4.8, 4.9, 2.8, 2.7 2.1, 3.1 2.9 4.0 2.3 2.3 1.3 5.2 2.4
Name of mine Kuroko-typedeposits Kunitomi Horobetsu Kagenosawa Toya-Takarada Minamishiraoi Minamishiraoi
K-Ar age (Ma) 12.6 12.3 14.2 14.0 13.6 13.6
Disseminated-type deposits Hakuryu 6.5 Date 5.2
variation. Homogenization temperatures for the precious vein-type, base metal vein-type and Kuroko deposits are 180-280°C, 200-250°C, and 180-250°C, respectively. Ishiyama et al. (1987) estimated hydrogen and oxygen isotopic compositions of ore fluid responsible for the base metal vein-type deposits in Jokoku-Katsuraoka area, southwestern part of the province. Estimated 3180 and 3D values range from -49%o to -88%o and from -11.6%o to +5.5%o, respectively. Hattori and Sakai (1979) analyzed inclusion fluids from the Chitose A u - A g vein-type deposits and indicated that 3D and 31SO for the ore fluids are -65%o to -75%o and -4.8%o to -6.2%o, respectively. Interpretation of the geological and geochemical characteristics of the three types of ore deposits are given below. It is inferred that in the northern part of the province submarine volcanic rocks are thick in the central zone, while at marginal zone it is thin and the Plio-Pteistocene subaerial volcanic rocks are exposed. The vein-type deposits occur widely in the province. The precious vein-type deposits occur in relatively young (Plio-Pleistocene) volcanic rocks, while large base metal vein-type deposits (e.g., Toyoha, Inakuraishi, Ohe) and Kuroko deposits (e.g., Kunitomi) occur in central zone where thick Miocene submarine volcanic rocks are distributed (Figs. 1.149 and 1.150). Small base metal vein-type deposits occur in Paleozoic rocks in the southern part. 3348 of the vein-type deposits hosted by sedimentary rocks of the basement are low (less than 0%o), reflecting low 3348 of country rocks. However, they are scarce in number. 334S of the base metal vein-type deposits and Kuroko deposits are relatively high (average value; +4%0 to +5%o) and most probably influenced by sulfate sulfur of seawater
Chapter 1
212
15
d- 10 0 Z .=_
5
O9
0
2OO
30O
Filling Temp. (~C) Figure 1.152. Salinity (NaCl eq.wt.%) and filling temperatures for the base metal vein-type (solid circle), Kuroko-type (solid square) and precious vein-type deposits (open circle) in southwest Hokkaido (Shikazono and Shimizu, 1993).
trapped in the Green tuff rocks. ~34S of the precious metal vein-type deposits closely associated with Plio-Pleistocene subaerial volcanic rocks have igneous value (ca. +2%o), suggesting that sulfide sulfur was extracted from subaeriat country rocks. The causes for the different site of hydrothermal activity (submarine and subaerial environment) could be considered in terms of tectonic and geologic evolution of this metallogenic province from middle Miocene to Pleistocene. During Miocene age most of this province was in a submarine environment. Violent submarine volcanism (bimodal and basic type) took place at Miocene age in this province. This geologic environment may be related to an extensional stress regime (Uyeda and Kanamori, 1979). The Kuroko deposits have been formed related to this tectonic situation. From late Miocene, uplift took place due to the collision of Pacific plate to North American plate under the Kuril arc and Japanese island arc. Watanabe (1986, 1989, 1990a,b, 1991) studied the vein pattern, the age of veintype deposits and the volcanic rocks in southwest Hokkaido and showed that the major veins such as those at the Toyoha and Chitose have been formed at dextral strike-slip movement of an E - W trend, and those veins are situated at the west-southwest extension of the maximum displaced zone within the dextral shear belt along the Kuril arc. Watanabe (1990b) also showed that the veins in the Sapporo-Iwanai district strike E - W and are oblique to the N W - S E volcanic chains which are sub-parallel to the maximum principal stress estimated in southwest Hokkaido during Late Miocene to Holocene and oblique subduction of Pacific Plate was active during the Plio-Pleistocene age. Yamagishi and Watanabe (1986) studied the geologic faults, the dykes, the veintype deposits and active faults in the rocks of the middle Miocene to Quaternary in
Miocene-Pliocene Hydrothermal Ore Deposits
213
this province and clarified that in the middle to late Miocene, southwest Hokkaido was characterized by a tensional stress field of an E - W to N W - S E direction, while in the late Miocene to Quaternary this area was in a compressional stress field. They showed that the main ore veins are recognized as sheared fractures arranged mostly in an E - W direction indicating the signal trajectory lines of an E - W to NW direction. It is noteworthy that the stress field of N W - S E direction estimated from a province of dikes during Miocene is parallel to the distribution trend of the Kuroko deposits in northeastern part of the district. Yamagishi and Watanabe (1986) suggest that the dip of subduction of Pacific plate during Miocene might have been steeper than during Plio-Pleistocene, considering the mode of subduction of Pacific plate in Northwest Honshu (Niitsuma, 1979). Usually active plate subduction with gentle dip causes uplift of land and expansion of land area (Uyeda and Kanamori, 1979). Such change from steep subduction to gentle one caused the changes in site of hydrothermal activity from submarine area during Miocene to subaerial area by Plio-Pleistocene to present. This change from extensional stress regime to compressional regime may correspond to the change of mode of subduction from Mariana-type to Chilean-type as defined by Uyeda and Kanamori (1979). It has been pointed out that this change has occurred at about 5 Ma in Northeast Honshu (Sugi et al. 1983; Ohmoto et al., 1983), and about 7 Ma in southwest Hokkaido (Otsuki, 1989). 1.5.4. Geochemical features of sedimentary rocks formed in the Japan Sea as a proxy for hydrothermal activity As noted already, intense submarine hydrothermal activity took place in the Japan Sea in 15-12 Ma, associated with Kuroko mineralization. However, it is uncertain that submarine hydrothermal activities associated with the Kuroko mineralization took place in the other periods from middle Miocene to present in the Japan Sea. Therefore, the geochemical features of sedimentary rocks which formed from the Japan Sea at these ages have been studied by the author because they are better indicator of age of hydrothermal activities than those of hydrothermally altered igneous rocks because the samples of continuous age of sedimentation are able to be collected and the ages are precisely determined based on microfossil data (foraminiferal, radioralian and diatom assemblages). Thick sedimentary pile from middle Miocene to late Pliocene is exposed in the Oga Peninsula, northern Honshu, Japan (Fig. 1.153). Age of the sedimentary rocks has been determined by microfossil data. Thus, the sedimentary rocks in the Oga Peninsula where type localities of Miocene sedimentary rocks in northern Japan are well exposed have been studied to elucidate the paleoenvironmental change of the Japan Sea (Watanabe et al., 1994a,b). Kimura (1998) obtained geochemical features of these rocks (isotopic and chemical compositions) and found that regional tectonics (uplift of Himalayan and Tibetan region) affect paleo-oceanic environment (oxidation-reduction condition, biogenic productivity). However, in their studies, no detailed discussions on the causes for the intensity and periodicity of hydrothermal activity, and temporal relationship between hydrothermal activity, volcanism and tectonics in the Japan Sea area were discussed. They considered only the time range from ca. 14 Ma to ca. 5 Ma.
214
Chapter 1
NYUDOZAKI
SARUgAWA
i
NISHIKUROSAWA ~
NOMI/RA
K1TAURA
ANDEN L"
' RECLAIMED
LAND
AIKAWA
TOGABAY~_I~j ~NINIOME.G ..A. .TAICHLNOMEG . ATA S A N ~ O M E -GATA
KAMO
Mr, K E N A S H I Y A M A JAPAN SEA
MONZEN
SUGOROKU
MINAMIH IRt..qAWA 0NNAGAWA
5kin
['~
~
Kampu-zanVOlCaNOs
Megata volcanic ejecta
~ KatanishiFormation ~ ~'/tJ Togapumicebed i ~rl~ ShibikawaFormation ~ ~
FunakawaFormation OnnagawaFormation NishikurosawaFormation DaijlmaFormation
WakimotoFormation ~
MonzenFormation
Kitaura Formation
AkashimaFomation
~
Figure 1.153. Geologic sketch map of the Oga peninsula.
Therefore, the wider time range from middle Miocene to present is considered below based on available age data on hydrothermal ore deposits (Kuroko deposits, epithermal vein deposits) and hydrothermal alteration in the mine areas in Northeast Japan.
Miocene-Pliocene Hydrothermal Ore Deposits
215
The geochemical features of the sedimentary rocks in the Oga Peninsula and the hydrothermal activity in Japan Sea deduced from these features are described below. The area is located at the west of Oga Peninsula (Fig. 1.153) and is composed of Miocene sedimentary rocks (Nishikurosawa, Onnagawa and Funakawa Formations). The Nishikurosawa Formation is composed of siltstone, mudstone, conglomerate and sandstone. Siltstone and mudstone contain foraminiferal fossil such as GIoborotalia birnageae, and G. denseconnexa, indicating Zone N. 9 by Blow (1969). The upper part is characterized by glauconite-bearing sedimentary rock. The total thickness is about 150 m. The Onnagawa Formation conformably overlies the Nishikurosawa Formation and is composed of siliceous shale and shale. The rocks are characterized by organicrich laminated diatomaceous deposits, siliceous microfossils, and fish bones, while foraminiferal fossil is poor in amounts. The total thickness is about 300 m. The age of base of the Onnagawa Formation is estimated to be 12.9 Ma based on diatoms (Koizumi and Matoba, 1989). The age of the top of the Formation is 5.8 Ma. The Funakawa Formation conformably overlies the Onnagawa Formation. The thickness is about 1000 m. The Formation is composed of siltstone, intercalated by tuff and tuffaceous siltstone. Foraminiferat fossil and siliceous microfossil exist in the formation. Average analytical data on the sedimentary rocks are compared with NASC (North American Shale Composite; Gromet et al., 1984) in Fig. 1.154. Average chemical compositions of the Nishikurosawa are very similar to NASC. However, Fe203, MgO, TiO2, MnO and P205 contents are higher and K20 content is lower than those of NASC, respectively. SiO2 and A1203 contents are similar with each other. The Onnagawa Formation is characterized by very high SiO2 content (average 89.12 wt%). Most of the other major element contents are lower than those of NASC, although MnO and P205 contents of the Onnagawa are higher. Average major element contents except SiO2 content of the Funakawa are very similar to NASC, but SiO2, Na20 and MnO contents are higher and the other element contents are lower than the NASC. Analytical data on minor element contents are compared with Average Shale by Turekian (1972) (Fig. 1.155). The average contents of most of the elements except Zn and Ba of the Nishikurosawa are same in orders to NASC. Average contents of minor elements of the Onnagawa are mostly lower than Average Shale. However, Cu, Zn, Mo, Ba and U contents are anomalously high compared with Average Shale. Average contents of minor elements of the Funakawa are lower than NASC which is due to slightly higher SiO2 content of the Funakawa. REE pattern normalized by NASC (REE contents; Goldstein and Jacobsen, 1988) are shown in Fig. 1.156. Nishikurosawa Formation is characterized by Eu positive anomaly which is calculated by the following equation: Eu/Eu* = (2Eu/Eu~ASC)/(Sm/SmNASC -k-Gd/GdNASC)
(1-75)
Chapter 1
216 (wI%}
I00
~A)
N/shik~t,
shale
rasa~;,~
8 ~g 40 NK¢-~eries s v e r a z e
2{) l{
-
@rome~. e~
~L 1984]
t :
.---
10E
(B) O n n a g a w a harod s h a l e
......~
OO-ser es @r~met
\
N
80
4O ~ver~e
et d
.
i9~
\
\
-
\ \
I
2
Miocene-Pliocene Hydrothermal Ore Deposits
217
4~ 'i } t{
k
Figure 1.154. Major element composition of studied rock samples. (A) Nishikurosawa shale, (B) Onnagawa hard shale, (C) Funakawa shale. Shaded area represents ranges.
The REE pattern for the Funakawa plot within 0.1-1 of NASC normalized value (Fig. 1.156). Ce and Eu anomaly for the Funakawa are weak (Ce/Ce* = 1.20 4-0.01; Eu/Eu* = 1.09 4- 0.05) (Fig. 1.157). The REE pattern for the Nishikurosawa exhibits slightly light REE enriched one (La/Yb = 0.55 4- 0.09) and plots close to the line of NASC normalized value -----1. Positive Ce anomaly is also found in the Nishikurosawa. This anomaly is defined by the following equation. Ce/Ce* = (3 Ce/CeNAS¢)/(2La/LaNASC + Nd/NdNAsC)
(1-76)
Hydrothermal solution venting from midocean ridges and back-arc basins has positive Eu anomaly (Klinkhammer et al., 1983; Michard et al., 1983; Mitra, 1994; Shikazono, 1999a) (Fig. 1.158). Therefore, the positive Eu anomaly of the sedimentary rocks is thought to be due to a contribution of hydrothermal solution. In order to know the contribution of hydrothermal solution the positive Eu anomaly of seawater (Eu/Eus*awater) is useful.
Chapter 1
218
' ppm~ f1
C
i
]ktrcB~ 1~;2]
i
I
[
~y °
~0~
m a h a ~ d sAale
fB~ 0 7 ~
i0/
° b" ......
%#
{ J
V
~n
I
,¢
{ h
~r ~e~
Zr y
?~ Nb
A ..........
Hf
R
Ba
Ph %~
U
Miocene Pliocene Hydrothermal Ore Deposits
219
(ppm) }£
[ ........ r - - - 7 - - 7 - - - ]
.............................................................
~ z t ;k~ a~
"
~
...........
192°9.
i
[ I} ~--
~
ot
Figure 1.155. Trace element composition of studied rock samples. (A) Nishikurosawa shale, (B) Onnagawa hard shale, (C) Funakawa shale. Shaded area represents ranges. The anomaly at that time is estimated based on the following equation. gu/gUr*oc k = n (gu/gut%rrigeneous) -Jr (1 - n)(gu/gus%awater )
(1-77)
where n is proportion of terrigeneous component in sediments, Eu/Eur*ck is Eu positive anomaly of sedimentary rocks, and Eu/Eut*rrigeneous is positive Eu anomaly of terrigeneous component. Eu/Eus*awater estimated based on the above equation is shown in Fig. 1.159. It is evident in this figure that intense hydrothermal activity occurred in the Nishikurosawa stage (14-13 Ma), it was very high " (Eu/Euseawater * -- 2.4) at the Nishikurosawa/Onnagawa boundary (12.6 Ma) which is similar to that of the sediments below the hydrothermal pool of Red Sea (Fish debris, Eu/Eu* = 2.4-4.9; Oudin and Cocherie, 1988), the intensity of hydrothermal activity decreased in the Onnagawa stage towards the Funakawa stage, and intense hydrothermal activity occurred intermittently at ca. 12.6, ca. 10.5 and ca. 8.2 Ma. High Zn, Cu and Ba contents of the Nishikurosawa Formation also indicate that the intense hydrothermal activity occurred at these ages. Magnetic susceptibility data are inferred to have been reflected by hydrothermal activity. The magnetic susceptibility data on the sedimentary rocks are shown in
220
Chapter 1 10
Funakawa shale
0.1
0,01
l
i
l
i
l
l
l
l
l
l
l
l
l
l
10 Onnagawa hardshale N
o~
@
=, 0.1
r/3
.< Z
l
l
l
f
i
l
i
l
l
l
l
I
10
0,01
Nishikurosawa shale
O.I
(1,(11
I
La Ce
I
t
I
I
v
NdSmEuGd
,
~
Dy
I
i
Er
I
E
YbLu
Figure 1.156. NASC-normalizedREE patterns. REE data of NASC obtained by Goldsteinand Jacobsen (1988) are used as normalizationvalue.
Fig. 1.160. The magnetic susceptibility values of the Nishikurosawa sedimentary rocks are high but those of the Onnagawa are relatively low. The values of the Funakawa are lower than those of the Nishikurosawa but higher than those of the Onnagawa. The values around the Nishikurosawa/Onnagawa boundary are the highest and they are variable in the Nishikurosawa and in the Onnagawa. The values of the Onnagawa are relatively high at 14-13 Ma, 12.6 Ma, 10.5 Ma, and 8.2 Ma. This variation correlates to Eu/Eus*awater
Miocene-Pliocene Hydrothermal Ore Deposits 1.4 1.3
221
m
Eu/Eu, 1 J:o
o
ml
|E
o o
E
1.2
0
o ®
m B
o
0
e! a o
1.1
o
o no a n o m a l y ~
1.0
o
o
O e
Pl )sitive negative
. . . . . . . . . . . . . . . . . . . . .
0.9 NK-shale
0t-t3
OG-hard
! .... l iFK-shale 01-04
shale 01-21
Figure 1.157. Eu/Eu* values of studied rock samples. I
w -E
0.01
I
I
I
1
I
Z
I
~ 13*N
-~
o
J
J
1984
-
13°N F982
m 2 1 * N 1981 ~.
~ \
O
o,oo :
(xlO-3t
\
_
_
#.°
-
/\
J
\
E
oc-
-
/
Cl~
0 0.0001
-
-
-
i
t
Ce
Nd
I
I
L
Sm Eu Gd
L
I
I
Oy
Er
Yb
Figure 1.158. Rare earth elements in vent fluids from the East Pacific Rise (redrawn from Michard and Albarede, 1986). The Sm values for the 1984 I3°N set are thought to be somewhat too high because of contaminationfrom the samplingsyringes.Data for basalt from 13°N East Pacific Rise are from Michard et at. (1983) (Scott, 1997).
variation, suggesting that high magnetic susceptibility may be due to high content of magnetite in the sedimentary rocks. Magnetite is considered to have formed biogenetically because magnetite is very fine-grained. If the grain size of magnetite is less than 0.03 Ixm, magnetic susceptibility of sediments containing magnetite is very high (Ioka and Yamazaki, 1994) and magnetite is thought to be of biogenic origin (Yamazaki et al.,
Chapter 1
222 (A)
(B)
- 2.0 " ~ ~ [ ' ~ t ~
......., ~
H ydrothermaI sediment
Seawater
d~tal
* - l.OI--n..... 0.8 I
Modern*
]
[
'v~"~O.~Q~Q " "~O'
aly
negative >12.9
., ,
12
,
11
,
10
,
9
"
% ,
8
7
5.8>
Middle Miocene to early Pliocene
J~l~r'. (Ma)
]
Figure 1.159. Eu/Eu* values of (A) modern sediment, hydrothermal solution and seawater and (B) midMiocene to early Pliocene Japan Sea (gu/Eus*awater, see in text). Modern data are from the Pacific ocean, except fish debris in the hydrothermal sediment (Red Sea). Modern* is defined here as including Quaternary time. Data sources: Hydrothermal solution (Michard and Albarede, 1986), seawater, fish debris in the Atlantic II deep (Oudin and Cocherie, 1988), hydrothermal sediment (Ruhlin and Owen, 1986).
1991). Probably, iron of biogenic magnetite was originated from hydrothermal solution. It is considered that ferric iron of hydrothermal solution was oxidized by iron oxidizing bacteria to form magnetite. The ages of Neogene mineralization and hydrothermal alteration in and around the Northeast Honshu and Hokkaido have been determined by K - A r data on K-minerals (K-feldspar, sericite). These data are summarized in Fig. 1.147 and Table 1.26. It seems clear by comparing Fig. 1.159 with Table 1.26 that the ages of hydrothermal mineralization and alterations determined by K - A r age dating are consistent with those of sedimentary rocks affected by hydrothermal activity in the Oga. Hydrothermal activities were intense at ca. 14-13 Ma, 12.6 Ma, 10.5 Ma, and 8.2 Ma. As already noted, major igneous and hydrothermal activities in the Japanese Islands seems likely to have taken place nearly at same times as the stress changes (Fig. 1.147). The stress field changes occurred at about 22 Ma, 15 Ma, 12 Ma and 8 Ma in the southern part of Northwest Japan and at about 15 Ma in the eastern part of Southwest Japan (Takeuchi, 1987) (Fig. 1.147). Kuroko and vein-type mineralization occurred at 23 Ma (Sado), 15-14 Ma (Kuroko in Hokuroku, Sado), 12.8 Ma (Karuizawa), 8 Ma (Takadama), 5 Ma (Hosokura), 3 Ma (Yatani) (Shikazono, 1985e, 1987b; Shikazono and Tsunakawa, 1982; Sugaki et al., 1986; Sawai and Itaya, 1996; Otsuki, 1989).
Miocene-Pliocene Hydrothermal Ore Deposits
223
(10 -3 SI) m!
1.0 t
Magnetic
2
susceptibility 0.8 0.3
J
Q m
0.2 [] WW
0.I
0 ® 0
0
0
®® @ ®
0 i
t
t
I
i
I
I
I
t
NK-shale 01-13
i
;
I+
I ~79
I ®I [ I I I I I I l0I T I I I I
OG-hard shale 01-21
0 fin Oi @ II FK-shale 01-04
Figure 1.160. Magnetic susceptibility value of studied rock samples. Most younger two samples (NK-12 and -13) of Nishikurosawa shale have anomalously high magnetic susceptibility value.
The changes in stress fields, and intensities of igneous and hydrothermal activities except 12.6 Ma seem correlate to oscillatory motion of the Pacific plate (Jackson's episodes; Jackson and Shaw, 1975; Jackson et al., 1975) (Masuda, 1984). Masuda (1984) and Takeuchi (1987) pointed out that the oscillatory motion of Pacific plate during the last 42 Ma correlates with magmatism, the intensity of tectonism, the change in stress field and the history of sedimentary basin in arc-trench system (Fig. 1.147). The above argument also suggests that the hydrothermal mineralizations in arc and back arc systems relate to the oscillatory motion of the Pacific plate. It is generally accepted that Kuroko deposits formed under the submarine environment, while polymetallic vein-type deposits in central and Northwest Japan (Ashio, Tsugu, Kishu, Obira, etc.) under the subaeriat environment. This spatial difference in the distribution of back-arc and vein-type deposits is found also for present-day mineralizations. For example, back-arc deposits are forming at Okinawa Trough, while base-metal and precious-metal precipitations are occurring on land such as at Ibusuki and Beppu geothermal area, Kyushu. Recent gold mineralizations (ca. 1 Ma) occurred at Noya and Hishikari areas in southern Kyushu. These mineralizations can be regarded as almost contemporaneous mineralization with the Okinawa mineralization. This kind of temporal and spatial relationship between epithermal Au vein-type mineralization and back-arc mineralization are found also in the Izu-Bonin area. Seafloor
Chapter 1
224 TABLE 1.26
Isotopic ages of metalliferous veins, altered country rocks and the igneous rocks which have genetical relations to the mineralization (Otsuki, 1989) Mine
Age (Ma) Mineral
Sanru Ohe Toyoha Inatoyo-Toyohiro Chitose Horobetsu Kagenosawa Koryu Hakuryu Date Ani Koaizawa Kameyamamori Arakawaohsawa Rata Sugisawa Kakkonda Hosokura Yatani
Takatama Karuizawa Ohizumi Sado
Ashio Nehazawa Chichibu
12.4 3.3 2.2 4.7 3.6 12.3 14.2 1.1 6.5 5.2 11 12.6 8.1 13.4 9.4
5.8 3.3 3.4 3.6 8.4 12.8
Metals
Vein strike
Au.Ag Mn>>Zn-Pb Zn > Pb>>Mn
N70°E, N85W; N15W N55°-85°W N40°-50°W, EW NS N65°E-N80W
Rock 6.6 8.4 8.5 3.4 10.3
12-15 23.6 15.2 8.7 13.6 12.0 4.9 9.7 9 5.2
5.3 13.4 14.5 22.1 24.4 14.8 5.0-5.7
Au-Ag Au-Ag-Cu.Pb.Zn Au.Ag.Cu-Pb.Zn Au-Ag Au.Ag.Cu.Pb.Zn Au-Ag-Cu-Pb.Zn Cu Ag.Cu.Pb.Zn Au.Ag.Cu.Pb-Zn Zn.Pb Au.Ag.Cu.Pb-Zn Au-Ag-Cu-Pb.Zn
N70°E, N55°-80°W N70° 90°E N80°-90°E ENE ENE N23°-70°E, N80°W N50°-60°W N25°W WNW-EW-NS N45°E, EW, NS NW N30°-80°E NNW-NW
Zn.Pb Zn.Pb
ENE, EW, NW EW, NE, NW
Pb.Zn Au-Ag Au.Ag Au.Ag Ag.(Pb-Zn) Zn.Pb.Cu Au.Ag
EW, 70W N70.W-EW-N45E
Cu Au-Ag 6.6
WNW, ENE, NNW N60°W, N60°E EW-ENE + NS + NW N65°-90°E, NW
N50°-70°E, EW N70 °, 85°E NS
b a s e - m e t a l m i n e r a l i z a t i o n s i m i l a r to the K u r o k o m i n e r a l i z a t i o n is f o u n d in I z u - B o n i n s e a f l o o r ( l i z a s a et al., 1999), w h e r e a s e p i t h e r m a l g o l d v e i n - t y p e m i n e r a l i z a t i o n o c c u r r e d at v e r y r e c e n t age (ca. 1 M a ) at Izu P e n i n s u l a , c e n t r a l Japan. T h i s spatial d i f f e r e n c e is c o n s i s t e n t w i t h the d i s t r i b u t i o n w i t h that o f h y d r o t h e r m a l d e p o s i t s o f m i d d l e M i o c e n e ( K u r o k o a n d p o l y m e t a l l i c v e i n - t y p e d e p o s i t s in Japan). I f the s u b m a r i n e vs. s u b a e r i a l h y p o t h e s i s m e n t i o n e d a b o v e is correct, it is e x p e c t e d t h a t K u r o k o d e p o s i t s are f o u n d in the m a r i n e s e d i m e n t a r y a n d v o l c a n i c h o r i z o n s w h o s e
Miocene-Pliocene Hydrothermal Ore Deposits
225
ages are 12.6 Ma, 10.5 Ma, 8.5 Ma and 5 Ma as well as 15-14 Ma (Kuroko deposits in Hokuroku area) and present-day (Okinawa and Izu-Ogasawara deposits). Large Kuroko deposits formed at 15-14 Ma (ore deposits in Hokuroku and Hokkaido) and at present (Okinawa Trough, Izu-Ogasawara). These ages correspond to rapid change of oscillatory motion of Pacific plate (Jackson's episodes). The change in the motion of the Pacific plate occurred at 15-14 Ma, 10.5 Ma, 8 Ma, 6 Ma, 3-2 Ma and probably present (Fig. 1.161). Thus, small Kuroko deposits probably formed at submarine environment at 10.5 Ma, 8 Ma, 6 Ma and 3-2 Ma, although these ore deposits have not been discovered. However, it is noteworthy that the 12.6 Ma corresponding to the Nishikurosawa/Onnagawa boundary when intense igneous and hydrothermal activities occurred does not correlate to Jackson's episode. The geochemical characteristics of the Nishikurosawa/Onnagawa boundary (12.6 Ma) are distinct from those of the other ages affected by hydrothermal activity (15-14 Ma, 10.5 Ma, 8.5 Ma); The positive Eu anomaly and magnetic susceptility were the highest compared with the other ages, glauconite is abundant, paleoenvironment was extremely reducing, and mass extinction occurred. This difference suggests that the cause for the volcanism and hydrothermal activity at that time were different from the other ages. The cause for hydrothermal and igneous activities at 12.6 Ma is uncertain because of no relationship with Jackson's episode. However, it is speculated that bolide impact may affect igneous and hydrothermal activities at this age. In addition to no relationship with Jackson's episode, this speculation comes from that the Earth's surface environmental change at the other geologic boundaries such as Eocene/Oligocene which is considered to have been influenced by bolide impact (Keller, 1986; Montanari et al., 1993; Vohof et al., 2000). Bolide impact may be important. But, further work should be clearly necessary to evaluate this hypothesis which is highly speculative.
1.5.5. Mode of subduetion and formation of back-arc basin
Uyeda and Kanamori (1979) divided mode of subduction into two types: Marianatype characterized steep subduction and Chilean-type characterized by gentle subduction estimated from the dip of Benioff-Wadati zones (Fig. 1.161). The geological phenomena associated with these subductions are shown in Fig. 1.162. It is inferred that the change in mode of subduction from Mariana-type to Chilean-type occurred at ca. 5 Ma in Northeast Honshu. Kuroko deposits are associated with Mariana-type, whereas epithermal vein-type deposits (particularly Au-Ag deposits) with Chilean-type. The above-mentioned view on the evolutionary history in metallogenic provinces from the formation of massive sulfide deposits at early-stage to that of the vein-type deposits in late-stage is consistent with the view by Cathles (1986) who has shown that massive sulfide deposits have been formed at an earlier stage than gold-quartz veins in Archean greenstone belt. Therefore, it is likely that the change in style of mineralization from massive sulfide deposits to gold vein-type deposits caused by the change in mode of subduction is common geologic phenomenon from Archean to Quaternary age in an Island arc-back-arc geologic setting. However, the relationship between the mode of
Chapter i
226 . Andaman M. America (12 N), C. Aleution, L. Antitles 100 km -
"~--'~
~
S.Chile(40%) Sumatra(5~S)
/S.Chile (30"S)/
I~~__~.~ru
~,laska,~.
'
~
New
~ ~
~.
j
(10°S) N.Chile(20°S)
Hebridesd'~ ~k'~~.~ "% ~
0
Depth, km 100 200
.yo.yu.p'%\'<,
'New Zealand
Ik ~
~
• Solomons
~ ~It~ , ~
Philippine-- - - ~ N e w Britain
~ ~
300
'~ '~
NEJalpan 400 ~-,,~Kurl~,
Java"
Mariana
'
'
700
Figure1.161.DipofBenioff-Wadatizones(UyedaandKanamori,1979). subduction and style of mineralization during the geologic history is not clear and should be studied in future. Horikoshi (1975a) proposed the hypothesis that Kuroko deposits formed at the time changing from extensional environment to compressional environment. He thought that metal-bearing fluids in the crust were pushed out by the compression. In contrast, Otsuki (1990) and Yamaji (1990) thought that back-arc rift activity and rapid subsidence occurred at ca. 15 Ma under the extensional environment. Bimodal volcanism was dominant at that time. Therefore, it is inferred that Kuroko formation took place at extensional environment, rather than compressional environment favored by Horikoshi (1975a); formation of Kuroko deposits may be related to the opening of the Japan Sea. However, the relationship between the opening of the Japan Sea and Kuroko mineralization is unclear, and there are many opinions on the genesis of the Japan Sea and timing of opening of the Japan Sea. For instance, Otofuji et al. (1987) and Otofuji (1996) showed that Southwest Japan Sea opened at middle Miocene very rapidly and widely based on paleomagnetic data. However, the rate of opening of the Japan Sea estimated by Otofuji et al. (1987) is too high (60 cm/year) and seems to be unreasonable. Therefore, many workers think that opening of the Japan Sea did not occur very rapidly and widely at middle Miocene age, but it occurred from earlier age (25 Ma) and opening occurred several times by the slower rate (Nakajima and Nakagawa, 1994). Twice openings are favored by several workers (Nakajima and Nakagawa, 1994). It is considered that there were several axes of the opening based on K - A r ages of volcanic rocks (Kobayashi, 1983).
Miocene-Pliocene Hydrothermal Ore Deposits
Ce_,
¢
-,,,
227
d
.,o*o
Mar~aria
"~~~ "
Figure 1.i62. Two kinds of subduction boundaries (Uyedaand Kanamori, 1979).
In contrast to Southwest Japan, opening of northeastern part of the Japan Sea is unclear, compared with the southwestern part of Japan. Tosha and Hamano (1988) made a paleomagnetic study of Tertiary rocks of Oga Peninsula (northern Honshu) and considered that a counterclockwise rotation of Northeast Japan with respect to eastern Asia took place between about 22 Ma and 15 Ma and the before the rotation, Northeast Japan was situated along the east coast of the Asian continent. Lallemand and Jolivet (1986) have interpreted the opening of the Japan Sea as a pull-apart basin between two right-lateral strike-slip fault zones; the Yagsan-Tsushima fault to the west and the Tartary-Hidaka shear zone to the east. Southward migration of Southwest Japan as a drawer between a right-lateral system to the west and a left-lateral one to the east along the Tanakura Tectonic Line (Otsuki and Ehiro, 1978) or a clockwise rotation of Southwest Japan about a pole located in the Tsushima Strait (Otofuji et al., 1987), or a combination of both (Kobayashi, 1983) have been proposed.
228
Chapter 1
Although many discussions on the origin of the Japan Sea have been carried out since the pioneer works by Tokuda (1927), and Terada (1934) who considered that the Japan Sea formed by southward migration from the Asian continent, the problem of the mechanism of the opening of the Japan Sea remains unsolved as mentioned above. Kuroko formation occurred at middle Miocene (15-16 Ma). But the opening of Southwest and Northeast Japan Sea occurred probably from the age earlier than this age. Therefore, it is thought that the beginning of the opening of the Japan Sea did not directly relate to the Kuroko formation. Cathles (1983a) proposed failed rift hypothesis to explain unusual features of the Green tuff region and Kuroko deposits. He mentioned that this hypothesis can account for the distribution of mine districts in the Green tuff region, the observed extensive and substantial premineralization subsidence and postmineralization uplift in the region, and its volcanic evolution. He further pointed out that the Green tuff belt of Japan is strikingly similar in geology, tectonics, and mineral deposits to Archean greenstone belts, suggesting that Archean belts may be failed rifts. As mentioned above, formation of back-arc basins and marginal seas may be important for the formation of Kuroko and vein-type deposits, although genetic relationship between Kuroko formation and opening of the Japan Sea is not clear. For example, Horikoshi (1977) insists that vein-type deposits in Northeast Hokkaido did not form without the opening of Ohotsuku back-arc basin. In section 2.3 and in Chapter 3, it is shown that the formation of back-arc basins take important role for the mineralization (back-arc deposits (Kuroko deposits), epithermal Au veins) and global geochemical cycle. Thus, it must be worth considering the formation mechanism of back-arc basins. The origin of the back-arc basins has been investigated considerably (Karig, 1971; Sleep and Toks6z, 1971; Uyeda and Kanamori, 1979; Tamaki and Honza, 1991; Uyeda, 1991; Tamaki, 1995) and various explanations for the origin have been proposed. Tamaki and Honza (1991) summarized the previously proposed models and the currently plausible models of back-arc spreading (Fig. 1.163). Model 1 shown in Fig. 1.163 is a slab-induced upwelling model. Upwelling generated along the down going slab-mantle boundary (Karig, 1971) or by secondary convection that is introduced by the slab (Sleep and Toks6z, 1971) causes back-arc formation (Karig, 1971 ). Model 2 is a plume injection model. Miyashiro (1986) showed that a northward migration of "hot region" (a large high-temperature region at a deeper level) governed the formation of back-arc basins in the Western Pacific. Tatsumi et al. (1990) examined the distributions of Cenozoic basalts and active rift systems in the Northeast China region and Quaternary subduction unrelated volcanism in the northwestern part of New Zealand (Fig. 1.164) and concluded the following: (1) Injection of the asthenosphere into the mantle wedge and the dam-up effect of the subducted slab explains the rifting process in the Japan Sea (Fig. 1.165). (2) Injection of the asthenosphere was associated with the eastward horizontal convective flow caused by the upwelling of asthenosphere beneath the northeastern China region (Fig. 1.165).
Miocene-PIiocene Hydrothermal Ore Deposits
229 , J / ~ - ""--I
I=IS!~!I, ! i I ~ k ~ I m tb mOdi~ £ 8 (le
Mode 1i Ssb4ii :d c@d pw@t gmsd~i!
XXXX
X
XX
M~de 4. ILIt,q!! £ iO@l£
m~d~l ? I~+I
M s I e 2 P!ume i i ~ t i ~ n msd@
!\
j y' /' / i ( i s
Figure 1.163. ModeIs of backarc spreading (Tamaki and Honza, 1991).
(3) The active back-arc extension in the Okinawa Trough and the Taupo Depression in New Zealand can be explained by the injection model. Otofuji (1996) proposed a "double door" opening mode with a fast spreading rate of 21 cm/year for the evolution of the Japan Sea, caused by the injection of asthenosphere with a low viscosity beneath the Japan Sea area. Model 3 is a plate kinematic model. The retreat of a back-arc plate forms a back-arc basin (Dewey, 1980). Model 4 is also a plate kinematic model. The retreat of a fore arc plate forms a back-arc basin. This model seems attractive. Jackson et al. (1975) found the periodicities of rotational motions of the Pacific plate. When the direction of the Pacific plate changed and obliquely subducted, the compressional force of oceanic plate to continental plate decreases. That means that the retreat of fore arc plate occurs.
Chapter 1
230
L NE CHINA UPWELLING
40N
L/ j~jl "~ • ".i I i ,
,
, °
°.
t~,,, e
:.,
•
B I":
140E
20N 120E Figure 1.164. Distribution of Cenozoic basalts and active rift systems in the northeast China region. Arrows indicate the horizontal convective current in the upper mantle associated with the upwelling of the asthenosphere beneath the region. A: Baikal Rift; B: Shanxi Graben; C: Tancheng-Lujiang Fault; D: Okinawa Trough (Tatsumi et aI., 1990).
INTRA-PLATE VOLCANISM CONTINENTAL RIFT SYSTEM P
+
+
÷
•
° E4
+'m,(÷
+'lrff+
l lll {11{
BACKARC BASIN
o,
;
÷
+
VOLCANIC BASIN
. .
°
5rrrvrrrrrrff+
÷
,
.4, y ~ .
~ /~/;~'0"0 Asthen°sPhi re.~q~ Asthenosphere
Upwellingof Asthenosphere ~
f
e,~,~
Figure 1.165. Schematic diagram showing the effect of upwelling of asthenosphere beneath a continent which is adjacent to an arc-trench system (Tatsumi et al., i990).
Miocene-Pliocene Hydrothermal Ore Deposits
231
In any model, back-arc basins form under the extensional stress regime and are associated by Mariana-type subduction by Uyeda and Kanamori (1979) rather than Chilean-type subduction. In this tectonic situation, intense bimodal volcanism and associated seawater circulation occur, resulting to the formation of Kuroko deposits on the seafloor and formation of vein-type mineralization under subaerial condition and intense hydrothermal and volcanic CO2 fluxes to ocean and atmosphere. Such fluxes affect the long-term environmental changes (see Chapter 4).
1.6. Other hydrothermal ore deposits 1.6.1. Polymetallic vein-type deposits The Honshu arc is divided into Northeast and Southwest Honshu by the Tanakura Tectonic Line or the Itoigawa-Shizuoka Line (Fig. 1.166). There are many geological and geochemical differences in these two geologic provinces. Neogene altered volcanic rocks (Green tuff) are widely distributed in Northeast Japan, whereas Cretaceous-Neogene granitic rocks and metamorphic rocks are in Southwest Japan. 87Sr/S6Sr ratio of volcanic rocks and granitic rocks (Shibata and Ishihara, 1979) (Fig. 1.166) and ~348 and sulfur content of Quaternary volcanic rocks (Ueda and Sakai, 1984) are different in two provinces (Fig. 1.167). As for the mineralization at the middle Miocene age in the Northeast Japan, Kuroko and epithermal base-metal veins have been formed. No enrichment of Sn and W is found in these deposits. In contrast, in Southwest Japan, polymetallic veins (so-called xenothermal-type deposits in the sense of Buddington (1935) or subvolcanic hydrothermal type in the sense of Cissartz (1928, 1965) and Schneiderhrhn (1941, 1955) occur. Examples of these deposits are Ashio, Tsugu, Kishu and Obira. All these vein-type deposits have formed at middle Miocene age in western part of Tanakura Tectonic Line under subaerial environment. In these deposits, many base-metal elements (Sn, W, Cu, Pb, Zn) and small amounts of Au and Ag are concentrated. These deposits are associated with felsic volcanic and plutonic rocks along the Median Tectonic Line (MTL) or south of MTL. According to Ishihara (1977), these granitic rocks are ilmenite-series while granitic rocks in Green tuff region are magnetite-series. 3 1.6.1.1. Ashio deposit The Ashio mine which is located in middle Honshu is one of the largest Cu ore producer in Japan until it was shut down in 1973 (Kanehira, 1991). The ore deposits
3Ilmenite-series and magnetite-series granitic rocks are defined as follows (Ishihara, 1977): the magnetiteseries and ilmenite-series granitic rocks are distinguished by the presence or absence, respectively, of magnetite in polished sections.
Chapter 1
232
HOKKAIDO
TANAKURA TECTONIC LINE ,"
0
0
o
o.,d KYUSHU
MEDIAN TECTONIC LINE
d
O U
I
300 KM ,
i
I
Figure 1.166. Distribution or initial 878r/86Sr (r i) for the Cretaceous-Paleogene plutonic rocks. Open squares: gabbros; open circles: granites. Numbers indicate the last two or three digits of ri values. Solid squares and solid circles imply the magnetite-series; open squares and open circles the ilmenite-series. (Shibata and Ishihara, 1979).
formed at middle Miocene age, associated with rhyolitic intrusive activity (Nakamura, 1970). The deposits occur in rhyolitic body as a funnel-shaped mass (Figs. 1.168 and 1.169). The deposits are characterized by conspicuous metal zoning and polymetallic mineralization. From the centre to margin of the mine district, the following zonings are recognized; S n - W - B i - C u zone, Cu-As-Zn zone, and Z n - P b - C u - A s zone (Nakamura, 1970). 3180 values of quartz are +12.8%o to 12.2%o (SMOW) (early-stage ( S n - W - B i Cu) quartz with cassiterite), +11.5%o to +11.4%o (quartz with stannite), +10.3%o to +8.8%0 (middle Cu-As-Zn stage), +9.4%e to +6.4%0 (late Z n - P b - C u - A s stage) (Shimazaki et al, 1993). If temperature is 300°C estimated from homogenization temperatures
Miocene-Pliocene Hydrothermal Ore Deposits 240
WEST JAPAN
420
6
70 60
233
Southwest Japan (SW) • RyukyuArc (RY)
A[
50
v High 834S rock (HI)
RY
* Alkaline (AL) ~" 4O
@
0..
~" 30 20
®
HI
10
+5
501
+10
EAST JAPAN
~ I M
30
20
0
-2
;H
HK
~'
_.~_~
0
+20
o Akita-Komagatake(AK) Hakone (HK) * Nasu zone (NA) , Chokai zone (CH) = Izu-Mariana(IM)
4O
(3.
+1'5
NA
+5
,
+10 ~34s (%,,)
+15
Figure 1.167. Sulfur content v s . 3 3 4 8 value for Quaternary volcanic rocks of Japan. Field bounded by solid lines show two volcanoes (AK and HK), two volcanic zones (NA and CH) in Northeast Japan, three volcanic belts (IM, SW and RY), alkaline rocks (AL) and volcanic rocks of unusually high 334S values (HI) in Ryukyu belt. Symbols surrounded by small circles show 334S values of Satsuma-Iwojima volcanic rocks in West Japan (Ueda and Sakai, 1984).
of fluid inclusion throughout all stages of mineralization, ~180 values of hydrothermal solution are early-stage: +4%° to +6%o; middle stage: +2%o to +3%0; and late-stage: - 1 % o to +2%0 (Shimazaki et al., 1995). This variation indicates that hydrothermal solution at early-stage was dominantly o f magmatic origin (or igneous origin), and a contribution of meteoric water component increased with the stage of mineralization. Main opaque minerals are chalcopyrite, cassiterite, stannite, arsenopyrite, bismuthinite, pyrrhotite and sphalerite. The FeS content of sphalerite is high (about 18 tool% FeS). Dominant gangue minerals are quartz, sericite, kaolinite and siderite. The Fe content of chlorite is very high (Fig. 1.170). All these data suggest low fs2 and f o 2 conditions.
Chapter 1
234 A V
V," V
Jv
V
V V
v
V
V
,
I1[
"","7"
\
/
II1
~°~°~
~_-'f.~" B
I: (Central) Sn-W-Bi-Cu zone I1: (Intermediate) Cu-As-Zn zone Ill: (Marginal) Zn-Pb-Cu-As zone
Figure 1.168. Geologic map of the Ashio mining area, showing the mineral zoning in the Ashio rhyolitic body and the location of the "Kajika" deposits (alter Ashio mine, partly revised) (Nakamura, 1970).
~34S values o f sulfides are close to 0%~ (Shimazaki, 1985), suggesting magmatic or igneous origin and no contribution of seawater sulfur. These values are quite different from those o f Green tuff sulfur (Kuroko and epithermal base-metal vein-type deposits; +2%~ to +7%o).
1.6.1.2. Tsugu deposit The Tsugu g o l d - a n t i m o n y deposit is located in the central part o f Honshu, Japan (Fig. 1.171). Geologically, the Tsugu deposit occurs in the Ryoke metamorphic terrane,
Miocene-Pliocene Hydrothermal Ore Deposits m
1200 A 1000 8O0 600 400
I [ ~~TM,. ~
•
235
\I
~
/
2OO
0t Rhyoliteweldedtuff ['L~
~
Rhyolite
S,a,e
~
OreVei. ' ai,k; depos,,s
lternationof sandstone and slate
Basalbreccia
H : Hotel vein,
lO00m
~Chert
[ q ~ " - I Rhyolite(lithoklite) dike. ~
500
J : Jimbovein, S : Shinseivein, Y : Yokomabuvein,
Figure 1.169. Geologic section along the line A B in Fig. 1.168 showing the mineral zoning in the Ashio rhyolitic body (Nakamura, 1970).
Mg4AI/ /
/"\ \
/ • ', /
/
L/'_ _~1.42 / . . . .
/',,,,,' Mg6
/"\ \
z13\
', ../.
/"k 06 l U ~ Fe(Mn)]4AI2
\,it/
/
"x,~.T12 - IlL/ -
\ 10_4/03\
\~.~/ i \
--$~--2_~
',, .,, ,,, ,,, ,,, ,,, ,,, ,,, \ [Fe(Mn)]6
Figure 1.170. Diagram showing the octahedraI composition of chlorites from the subvolcanic hydrothermal deposits, propylite, and Kuroko deposits in Japan (Nakamura, 1970). Chlorite occurring as a gangue mineral in the subvolcanic hydrothermal deposits: Nos. 1, 2, 3 and 4: Chlorite from the Ashio copper mine. Nos. 5, 6, and 7: Chlorite from the Kishu mine. No. 8: Chlorite from the Arakawa mine. Nos. 9 and 10: Chlorite from the Ani mine. No. 11: Chlorite from the Osarizawa mine. Chlorite from the so-called propylite: No. 12: Chlorite from the Yugashima mine. No. 13: Chlorite from the Budo mine. Chlorite from the Kuroko deposits: No. 14: Chlorite from the Wanibuchi mine.
adjacent to the outer z o n e o f southwestern Japan. In this z o n e and in the R y o k e m e t a m o r phic terrane, A u deposits are rare, w h e r e a s there are m a n y H g - S b deposits. O n the other hand, as already noted, in the Tertiary s u b m a r i n e v o l c a n i c r e g i o n ( G r e e n tuff region) and Quaternary v o l c a n i c region, m a n y e p i t h e r m a l v e i n - t y p e deposits occur. M i n e s w h i c h h a v e p r o d u c e d b o t h Sb and A u are very rare. T h e r e f o r e , the T s u g u deposit is an unusual type o f A u deposit in Japan. I g n e o u s activity related to the m i n e r a l i z a t i o n indicates that the age o f m i n e r a l i z a t i o n was m i d d l e M i o c e n e and this age is different f r o m that o f e p i t h e r m a l A u - A g mineralization, m o s t l y late M i o c e n e to Pleistocene, as already noted.
Chapter 1
236
17= z,0o - - -
35 °
~.
Tsugu D e p o s i t
i 133°
0
200kin
138°
Figure 1.171. Map showing location of the Tsugu gold-antimony deposit. I: Green-tuff region, 2: outer zone of southwestern Japan, TTL: Tanakura tectonic line, 1STL: Itoigawa-Shizuoka tectonic line, MTL: Median tectonic line (Shikazono and Shimizu, 1988a).
The previous studies on the Tsugu deposit (Tsuboya, 1936; Tatsumi, 1948; Shikazono and Shimizu, 1988b) demonstrated that: (1) Opaque minerals include stibnite, jamesonite, cinnabar, gold, pyrite, pyrrhotite, arsenopyrite, marcasite, sphalerite, galena and chalcopyrite. (2) The ore minerals display a zonal distribution: gold and cinnabar are enriched in the upper part of the veins, and sphalerite, galena and chalcopyrite are more abundant in the deeper parts. Pyrrhotite and arsenopyrite are distributed throughout the veins. (3) From the mode of occurrence of opaque minerals it is considered that pyrrhotite and sphalerite were precipitated at an early-stage, gold, pyrite, marcasite, stibnite and cinnabar were precipitated at a late-stage, and arsenopyrite was precipitated throughout the mineralization period. (4) The accessory minerals are tourmaline and apatite. Hg content of electrum ranges from nil to 8.5 wt%, and Ag content ranges from
237
Miocene-Pliocene Hydrothermal Ore Deposits Au
(z-solid / so, t,o
j
•
•
•
Ag
Hg 10
20
30
40
50
60
70
80
90
Hg atomic % Figure 1.172. Chemical composition of mercurian gold plotted on a Au-Ag-Hg triangular diagram. Data sources: Tsugu (open square; Shikazono and Shimizu, 1988b); L~ngsele (solid circle; Nysten, 1986); Aitik, northern Sweden (solid triangle; Nysten, 1986); Kazakhstan (cross; Naz'mova and Spiridonov, 1979); Bajupura-Dariba (solid square; Basu et al., 1981)• Phase boundaries (solid lines) at 450°C were drawn from Basu et al. (1981) (Shikazono and Shimizu, 1988b).
6.0 to t 1.5 wt%. The chemical compositions of gold from the Tsugu deposit are plotted in terms of Au, Ag and Hg, together with data from the literature (Fig. 1.172). The Au/Ag ratios of electrum from the Tsugu deposit are high compared with those from the other mercurian gold occurrences. Note that Hg content varies inversely with Au content (Fig. 1.173), implying that Hg substitutes for Au. A frequency histogram of the Ag content of electrum from epithermal gold-silver vein-type deposits and the Tsugu deposit (Fig. 1.174) clearly indicates that the Au/Ag of electrum from the Tsugu deposit is higher than that from epithermal vein-type deposits. Electrum and cinnabar occur in the same part of the vein, suggesting that these minerals were in equilibrium. The following reaction can be used to determine which factors are important for controlling mercury content of electrum in equilibrium with cinnabar: (Hg) + 1/2S2 = HgS
(1-78)
where (Hg) denotes the Hg component in electrum. From the equilibrium relation for reaction (1-78), we obtain, NHg
1/(gl-78fs2 YHg) 1/2
(1-79)
where NHg is the atomic fraction of the Hg component in electrum, K1-78 is the equilibrium constant of reaction (1-78), i s 2 is the sulfur fugacity, and gHg is the activity coefficient of the Hg component in electrum.
238
Chapter 1 8
6
E
4
-r
2
0
70
75
80
85
90
Au atomic % Figure 1.173. Relationship between mercury and gold contents of mercurian gold from the Tsugu goldantimony deposit (Shikazono and Shimizu, 1988b).
From (1-79) it is clear that the Hg content of electrum is related to f s 2 and K1-78. Using the thermochemical data for reaction (1-78) by Barton and Skinner (1979), isoactivity lines for Hg in electrum may be drawn on a log fsz-temperature diagram (Fig. 1.175). At a given temperature, the activity of Hg in electrum increases with a decrease in fs2. Therefore, the f s 2 for the mercurian gold of the Tsugu deposit is inferred to be relatively low. Microprobe analyses of sphalerite indicate that sphalerite from the Tsugu deposit is markedly higher in iron than that from the epithermal Au-Ag vein-type deposits in Japan. Measurements of homogenization temperatures of fluid inclusions that were obtained from quartz coexisting with sphalerite, pyrrhotite and chalcopyrite of the earlystage of sulfide mineralization, and from quartz coexisting with electrum, arsenopyrite, pyrite and sericite of the late-stage of mineralization. Homogenization temperatures of fluid inclusions are 281-345°C for the early sulfide stage, and 275-295°C for the latestage of gold mineralization. These fluid inclusion data, stability fields of arsenopyrite, pyrite and pyrrhotite, and the iron content of sphalerite can define the possible ranges of sulfur fugacity and temperature for the Tsugu mineralization (Fig. I. 176). The ranges of is2 and temperature for epithermal Au-Ag vein-type deposits in Japan have been clearly defined based on the chemical composition of sphalerite and electrum, and homogenization temperatures of fluid inclusions (Shikazono, 1985d). Values of fs2 for the Tsugu deposit are lower than the typical ranges of values for the epithermal Au-Ag vein-type deposits in Japan (Fig. 1.176). Such a low i s 2 is in accord with the high Hg content of electrum in the Tsugu deposit.
Miocene-Pliocene Hydrothermal Ore Deposits
239
70 ¸
60
50 O ©
40
(1) ~4 30
20
10
20
40
60
Ag a t o m i c
80
I00
%
Figure 1.174. Frequency (number of analyses) histogram for Ag (atomic %) of gold from the Tsugu deposit (solid) and epithermal gold-silver vein-type deposits in Japan (open). I: sampte I; II: sample II. Data sources: Tsugu deposit (Shikazono and Shimizu, 1988b); epithermal gold-silver vein-type deposits in Japan (Shikazono 198l, 1986; Shikazono and Shimizu, 1988b).
~xx%
o .a -15
150
200
250
300 Temp.(°C)
Figure 1.175. Activity of S2-temperature diagram showing iso-Hg contents contours for gold. The calculations were carried out using thermochemical data of Craig and Barton (1973).
Chapter 1
240
/-
~
-10
'
J
'
o -
/
c
//////" ,
,
,
r
150
,
,
,
,
I
200
~
I
250
I
i
t
i
I
300
.
e
%t-.
,. ~,.o-
//...~,%.-'//
O k/~
J
i
-1 ~
i
-
i
i
i
i
1
350 Temp,(°C)
Figure 1.176. Possible ranges of activity of $2 (as2) and temperature for the Tsugu deposit (solid and dotted areas) and epithermal gold-silver vein-type deposits in Japan (shaded area) (Shikazono and Shimizu, 1988a).
Mercurian gold has been reported from several Au deposits: Longsele, Aitik, Kazakhstan, and Bajpura-Dariba (Naz'mova and Spiridonov, 19797 Basu et al., 1981, Nysten, 1986). The nature of the opaque minerals coexisting with mercurial gold in these minerals are pyrrhotite, galena, sphalerite, stibnite, and jamesonite, and less common are cubanite, arsenopyrite, pyrite, and owyheeite. Pyrrhotite, cubanite, and arsenopyrite suggest relatively low fs2, but the temperatures of formation of these deposits have not been estimated. 1.6.1.3. Kishu deposit The Kishu deposit, located at middle Honshu, is associated with Kumano acidic rocks intruded into the Shimanto and Kumano formations. K-Ar age determination indicates that the activity of Kumano acidic rocks occurred at middle Miocene (144-1 m.y.) (Ikebe, 1973). The stage of ore mineralization from early to late is (1) Cu-Fe (chalcopyrite, pyrite, chlorite, quartz), (2) Cu-Zn-Pb (chalocopyrite, sphalerite, galena, sericite, quartz, siderite, calcite, fluorite, chlorite), (3) Au-Ag (native gold, argentite, pyrite, sphalerite, chalcopyrite, chlorite, quartz, adularia, sericite), (4) calcite (Nakamura and Miyahisa, 1976). ~348 of sulfides from Kishu is very low (-14.2%o) (Shimazaki, 1985), indicating sedimentary sulfide sulfur source. 1.6.1.4. Obira deposit Obira deposit is located at south Kyushu. It occurs in Shimanto Group rocks composed of shale, and limestone, associated with middle Miocene granitic intrusive rocks. Polyascendant zoning and mineralization are observed in the Obira deposit. From early- to late-stage mineralizations are:
Miocene-Pliocene Hydrothermal Ore Deposits
241
cassiterite-apatite-tourmaline --+ cassiterite-wolframite-molybdenite-pyrite --+ pyrite-arsen•pyrite-pyrrh•tite-magnetite-cha•c•pyrite-stannite-bismuthinite-ga•enatetrahedrite --+ marcasite-pyrite --+ marcasite-Sb minerals (stibnite, jamesonite, berthierite, boulangerite). Occurrence of pyrrhotite, arsenopyrite and high iron content of sphalerite indicate low i s 2 and / O 2 conditions. Decrepitation temperatures of fluid inclusions in quartz from S n - C u - A s stage are 403-322°C (Enjoji and Takenouchi, 1976). ~34S values of sulfides are low (-6.7%o) (Shimazaki, 1985), suggesting a contribution of sedimentary sulfide sulfur. Distinct features of the above-mentioned polymetallic deposits are summarized as follows: (1) The FeS content of sphalerite and Fe content of chlorite from these deposits are high. (2) Temperature of formations is high and variable. (3) fs2 and f02 are low. (4) Many metals (Cu, Pb, Zn, Au, Ag, Sn, W, Bi) have been produced. (5) Orebody and regional metal zonings are conspicuous. (6) Salinity of inclusion fluids varies widely from high to low. (7) Common gangue minerals are kaolinite and sericite, but K-feldspar is not found, suggesting low pH of ore fluids. (8) The deposits have been formed under the subaerial environment. No contribution of seawater and low fs2 and f02 conditions are consistent with the geologic environment.
1.6.1.5. Temperature and sulfur fugacity estimated from iron and zinc partitioning between coexisting stannite and sphalerite and coexisting stannoidite and sphalerite Stannite is the most common tin sulfide mineral in the ore deposits associated with tin mineralization. This mineral sometimes contains appreciable amounts of zinc, together with iron. Several workers have suggested that the zinc and iron contents of stannite are related to temperature. With respect to the study of the phase relationships in the pseudobinary stannite-kesterite system, Springer (1972) proposed zincic stannite as a possible geothermometer mainly based on the chemical compositions of the two exsolved phases (stannite and kesterite). Nekrasov et al. (1979) and Nakamura and Shima (1982) experimentally determined the temperature dependency of iron and zinc partitioning between stannite and sphalerite. Because natural stannite contains a considerable amount of zinc, sphalerite contains a considerable amount of iron, and these contents can be easily analyzed using an electron microprobe, a stannite-sphalerite pair is expected to be a useful indicator of formation temperature and sulfur fugacity. Iron and zinc partitioning between stannite and sphalerite is represented by the exchange reaction. Cu2FeSnS4 (in stannite)+ ZnS (in sphalerite) = Cu2ZnSnS4 (in stannite)+ FeS (in sphalerite)
(1-80)
Chapter 1
242
"" 1.0 cn o
I .
~.
1
Nekrasov el 0t.0979),
looKd =l.27&.lOd'T-1-l.174
7 Nokomuro & Shimo ( 1982 ), log Kcl = ZS.103.T- 1_3fi
250
2.0
300
350
l~a
400 450
1.6
590~C
~4
f.2 103"T -1
Figure 1.177. Comparison between the stannite-sphalerite geothermometer after Nekrasov et al. (1979) and one after Nakamuraand Shima (1982). Crossbars indicate experimental uncertainties (Shimizu and Shikazono, 1985). 153
The partition coefficient, Kd, for the above reaction is expressed by,
Kd = (Xcu2FeSns4/Xcu2znsnS4)stannite/(XFeS/XZnS)sphalerite
(1-81)
where X denotes mole fraction of a given component in stannite and sphalerite. Nekrasov et al. (1979) and Nakamura and Shima (1982) reported a temperature dependency of iron and zinc partitioning between stannite and sphalerite (Fig. 1.177) logKd = 1.274 x 103T -I - 1.174 (Nekrasov et al., 1979)
(1-82)
logKd = 2.8103T -1 - 3.5 (Nakamura and Shima, 1982)
(1-83)
Both geothermometers are in agreement being close 380°C (Fig. 1.177). However, at the lower and higher temperatures the difference between the temperatures estimated from equations (1-82) and (1-83) becomes larger. Figure 1.178 represents a comparison between the stannite-sphalerite temperatures and homogenization temperatures of fluid inclusions or sulfur isotope temperatures. It can be seen in Fig. 1.178 that Nakamura and Shima's geothermometer would be rather consistent with the temperature estimated based on the fluid inclusions or sulfur isotope studies. It is notable that almost all stannite-sphalerite temperatures are within 30°C of average homogenization temperatures and sulfur isotope temperatures. Temperatures estimated based on equations (1-82 and 1-83) range from ca. 250°C to 350°C for both skarn-type and vein-type deposits except the Yatani epithermal Au-Ag vein-type deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
243
°C
°C
400.
400 /
E O
IoT ~K
or'
E 300,
(N
tO
200
200
100
100
@
C O D
"6 2
Nekrasov et al (1979) 0
Nakamura and Shima (1982) t
0 0
100
200
300 °C
100
200
i
300 °C
stannite-sphalerite geothermomentry Figure 1.178. Comparison between the stannite-sphalerite temperatures and fiIIing temperatures of fluid inclusions or sulfur isotope temperatures.NT: Nakatatsu, OB: Obira, KN: Kant, KG: Kuga, TM: Tsumo, KM: Kamioka, OT: Ohtani, KU: Kaneuchi, Ak: Akenobe, TT: Takatori,YT: Yatani (Shimizu and Shikazono, 1985).
The chemical compositions of the coexisting stannite and sphalerite are plotted on log(XFeS/XZnS)sphaterite--log(Xcu2Fe2SnSg/XCu2ZnSnS4)stannite diagrams (Figs. 1.179 and 1.t80). Iso-logfs2 lines on these diagrams are based on equation (1-83), and thermochemical data on the F e - Z n - S system by Scott and Barnes (1971). The line 20.8 mol% FeS in sphalerite, which corresponds to the composition of sphalerite in equilibrium with pyrite and pyrrhotite at 1 bar as determined by Boorman (1967), is also given on the diagrams. It is evident in Figs. 1.179 and 1.180 that the FeS content of sphalerite coexisting with pyrite is generally lower than that of sphalerite coexisting with pyrrhotite or both pyrite and pyrrhotite. Iso-FeS content lines for sphalerite in equilibrium with pyrite or pyrrhotite were drawn on the log fs2-temperature diagram (Figs. 1.179 and 1.180) using thermochemical data by Scott and Barnes (1971) and Barton and Skinner (1979). The relationship between the iron content of stannite in equilibrium with sphalerite and pyrite or with sphalerite and pyrrhotite was derived based on thermochemical data by Scott and Barnes (1971), Barton and Skinner (1979) and Nakamura and Shima (1982). These types of deposits are skarn-type polymetallic (Sn, W, Cu, Zn, Pb, Au, Ag) vein-type and S n - W vein-type deposits. As shown in Fig. 1.181, the fs2-temperature range for each type of deposits is different; at a given temperature, fs2 increases from S n - W vein-type through skarn-type to polymetallic vein-type deposits. It is interesting to note
Chapter 1
244
250°C
¥r
].
Ti.0
am
/ ~ 1 2 C U')
u7 rt/~
U
(D
450°C
0
0
-t
-t
0
log XFeS in spholerite XZnS Figure 1.179. log(XFeS/XZnS)sphalerite--Iog(XCuzFeSnS4/Xcu2ZnSnN4)stannite diagram showing that sphalerite and stannite are associated with pyrrhotite (Po) and/or pyrite (Py). Temperature lines are based on data by Nakamura and Shima (I982). Solid curves show logfsz based on data by Scott and Barnes (1971) in the pyrrhotite field. Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).
that i s 2 associated with gold-silver mineralization in polymetallic deposits is relatively high compared with those of the other types of deposits. This difference is similar to that found in epithermal vein-type deposits; fs2 of epithermal A u - A g veins is higher than that of base metal veins. Shimizu and Shikazono (1987) studied the compositional relations of coexisting stannoidite, sphalerite and tennantite-tetrahedrite (Fig. 1.182). Based on these data they estimated the sulfur fugacity of stannoidite-bearing tin ore. Considering the complementary work on stannite-bearing tin ores from Japanese ore deposits (Shimizu and Shikazono, 1985), a comparison between environmental conditions of these two types of tin sulfides was made. Their study is described below. The chemical compositions of coexisting sphalerite and tennantite-tetrahedrite from the mines were determined. Except the Ashio polymetallic deposits, the other deposits have been formed at late Cretaceous related to felsic magmatism. Scanning patterns and many analytical data obtained by electron-microprobe analyzer reveal that most stannoidite grains are compositionally homogeneous. There is
Miocene-Pliocene Hydrothermal Ore Deposits
245
YT
250oc 9'/', ,
/iK@iL,
I C 0
/
3sooc
,..5ooc
C
o
,
/
-1 -2
.
Py
1
tog XFeS
in sphalerite
XZnS Figure 1.180. log(XFeS/XZnS)sphalerite Iog(Xcu2FeSnS4/XCu2ZnSnS4)stannitediagram showing each deposit is skarn-type or vein-type. Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).
a wide range in extent of Fe and Zn substitution in stannoidite (NFe/l~Zn between 2.03 and 14.4). The stannoidite from the Tada deposit is the richest in Zn (5.24 wt%, Zn, 9.09 wt% Fe) and has the approximate formula Cu8Fe2ZnSn2S12; that from the Konjo deposit is the richest in Fe (11.90 wt% Fe, 1.69 wt% Zn) and has the formula Cu8Fe2(Fe0.77Zn0.19)Sn2S12. From the Fe2+/Zn 2+ ratio of stannoidite, a continuous solid solution is inferred to exist between Cu8Fe2ZnSn2S12 and CusFe2FeSn2SI2. Iron and manganese contents of sphalerite coexisting with stannoidite are in the range from 0.17 to 1.93 wt%, and from 0.02 to 0.16 wt%, respectively. Lee et al. (1974) conducted an experimental study on the equilibrium for the assemblage of stannoidite-chalcopyrite-bornite-mawsonite-S2 (gas) in a temperature range from 430 to 300°C. Curves A and B in Fig. 1.183 correspond to fsz-temperature relationships for this equilibrium assemblage for aFe = 1 and aFe = 0.1, where aFe is the activity of the Cu8Fe2FeSn2S12 component in stannoidite solid solution. As mentioned already, Shimizu and Shikazono (1985) have estimated the is 2temperature range for stannite-bearing assemblages from Japanese vein-type and skarntype tin deposits. This estimated fs2-temperature region is also shown in Fig. 1.183. The fsz-temperature range for the formation of these two types of tin sulfides is different.
246
Chapter 1
Cn 0 I
10
Temperoture, *C 11
250 30O 35O *C Figure 1.181. Temperature-log fs: diagram. Skarn-type deposits (solid squares) are considered to be formed under lower fs2 condition than vein-type deposits (open squares). Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).
Microscopic observation suggests that stannite formed earlier than stannoidite on the scale of one polished section, although in general coexistence of stannoidite and stannite cannot be observed in the same polished section. If stannite formed at an earlier stage than stannoidite, it could be inferred that fs2 increased or that temperature decreased (or both) with evolution of tin mineralization. The Fe2+/Zn 2+ ratio of coexisting stannoidite, sphalerite and tennantite-tetrahedrite from the Tada, Omodani and Ohmidani deposits is low, compared with that from the other deposits such as the Ashio, Akenobe and Ikuno deposits (Fig. 1.182). The Tada, Omodani and Ohmidani deposits are characterized by polymetallic (Zn-Cu-Ag-Au; Zndominated) mineralization, and tungsten is not recovered from these deposits. On the other hand, the Akenobe, Ikuno and Ashio deposits are characterized by polymetallic (Cu-AnPb-Sn-W-Ag-Au-Bi) mineralization, and tungsten is recovered from these deposits. The Fe/Zn ratio of coexisting stannoidite, sphalerite and tennantite from the tungsten-bearing polymetallic deposits is higher than in those from the tungsten-free polymetallic deposits. High content of indium in chalcopyrite from tungsten-beating polymetallic deposits probably indicates a high temperature of formation. Sphalerite rarely appears where the Fe2+/Zn 2+ ratio of the coexisting stannoidite exceeds approximately 1. These differences imply that the tungsten-bearing deposits formed under lower fs2 or higher temperature conditions (or both) than tungsten-free deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
247
KONJO
2-
2
I~KUNO A ¢N
ASHIO
1
~OMO_~I
ASHIO N'ENOE
AKENOBEI
~,
/ FUKOKU
=~"KOKU
tOMO' ~DANI OHMIDANI 0
I 2 3 4 ( Fe 2" I Zn 2" )sp x 10 2
0
1 2 3 ( F e z. / Zn 2" )tenn-tetr
Fig 1.182. Fe2+/Zn2+ of sphalerite (left) and of tennantite-tetrahedrite-series mineral (right) as a function of FeZ+/Zn2+ in stannoidite (atomic proportions) (Shimizu and Shikazono, 1987).
Figure 1.183 shows i s 2 and temperature region for stannoidite-bearing tin deposits is different from that for stannite-bearing tin deposits.
1.6.2. Hg and Sb deposits Hg deposits are distributed along the MTL (Median-Tectonic Line) in Southwest Japan and in Northeast Hokkaido (Fig. 1.184). The deposits are vein or disseminated in form. The deposits are hosted by sedimentary and igneous rocks. No K - A r age data on the deposits are available. However, from the age of host rocks, the age of igneous activities along MTL and the studies on the movement of MTL and K - A r ages of Au-veins associated with Hg mineralization in Northeast Hokkaido (e.g., Khonomai), it is likely that these deposits formed at middle Miocene age. However, mercury mineralization in Kitami Province (north Hokkaido) occurred at approximately the same age as the epithermal gold-silver mineralization in the same district (4.5-5.3 Ma) (Maeda, 1997). Main ore minerals are few. They are cinnabar, metacinnabar, stibnite and pyrite. Gangue minerals are quartz and calcite. Fluid inclusion decrepitation temperatures
248
Chapter 1
0001
Stcmnoidite r e g i o n / 0005
~'-. . .
0
jf
10 ," /// B
Stotnnite r e g i o n (Shimizu & 5hikazono 1985)
./
15
A
"
Temperature 20
200
250
3 0
350"C
Figure 1.i83. Temperature-log fs2 diagram for stannoidite-bearing ores from Japanese veto-typedeposits. MW = Mawsonite, Sd = Stannoidite, Bn = Bornite, Cp = Chalcopyrite(Shimizu and Shikazono, 1987).
for quartz and calcite from the Itomuka Hg deposit are 328-337°C and 201-242°C, respectively (Enjoji and Takenouchi, 1976). ~34S of cinnabar and stibnite are wide in a range from -16%e to +15%o, and from -15%o to +4%0 for cinnabar and metacinnabar, and stibnite, respectively (Shikazono and Shimizu, unpublished). Usually, 334S values of cinnabar and stibnite are low less than 0%o except for these minerals from Green tuff region. These low 334S values imply that sulfide sulfur of Hg and Sb deposits came dominantly from sedimentary sulfur, and no contribution of seawater sulfur. The sulfur isotopic data are consistent with geologic environments of Hg and Sb deposits; Sedimentary rocks are dominant and marine rocks are not present in Sb-Hg mineralization districts. However, a few samples of stibnite and cinnabar from the deposits in Green tuff region display high ~34S values. In contrast of this interpretation on the origin of sulfur, Ishihara and Sasaki (1994) thought that sulfur came from ilmenite-series granitic rocks. However, these rocks are not found in the north Hokkaido. Horikoshi (1995) showed areal extension of volcanic belts since middle Miocene in the Japanese Islands. He suggested that Hg and Sb mineralization in outer zone of Southwest Japan and north Hokkaido (Kitami district) related to forearc igneous activities in trench side. ~34S values become heavier across the Northeast Japan arc from the trench to the back-arc side. He thought that the changes in tectonic character from forearc to arc environment controlled the ~34S values.
Miocene-Pliocene Hydrothermal Ore Deposits
249
p R
i Sado
O0
,
Sanin Kyoto s
g
% o okue~_ ~
~o~s p~o.4\~ce
e~ 're'=
\\ra
zone,
0 I
I
I
3 0 0 km I
qxl'l ,~*~"
o~te~
,0
• e • e o
Hg Ore Deposits Sb Ore Deposits Au-Ag Ore Deposits Pb-Zn Ore Deposits Sn-W Ore Deposits
o Vein o
m Skarn
D
Figure 1.184. Distribution of major sulfide deposits in Miocene volcanic and plutonic provinces. Class i and 2 are ore deposits (that is, the largest and second largest are shown together with the larger symbols). Class 3 is shown as a smaller circle and box. Tin ores occur associated with base-metal deposits at Mt. Okue, and tungsten deposits are concentrated in Yakushima, Kyushu (Ishihara and Sasaki, 1991).
1.6.3. Gold-quartz vein-type deposits (mesothermal-hypothermal vein-type deposits) 1.6.3.1. Geology, mineralogy and geochemistry G o l d - q u a r t z v e i n - t y p e ( m e s o t h e r m a l - t y p e a n d h y p o t h e r m a l - t y p e d e p o s i t s in t h e s e n s e o f L i n d g r e n ( 1 9 2 8 ) ) o c c u r in s e d i m e n t a r y t e r r a n e a s s o c i a t e d w i t h C r e t a c e o u s f e l s i c
Chapter 1
250
0
100
200km i
40° Nishizawa.~°h °
k
/ Hayochine ~Z~;i!~v//Owashi t - - - % ~ ~//Nojiri u ~ Shishiori
K i n k e i \ F::~::/\ ~ ~ i ' )
x AikGwa hio owo
o Amo---------~E~r,', ~ ~ _ 35°
~3 0
_/ .....
133°
/__..~
~~b'~--.
:~; \ ~
~
Suwa
AShiyasu
138°
Figure 1.185. Locations of hypo/mesothermal vein-type deposits in Japan. Abbreviations are the same as those in Fig. 1.62. Open circle: hypo/mesothermal Au vein-type deposits. Solid circle: hypo/mesothermal po]ymetallic vein-type deposits (Shikazono and Shimizu, 1988a).
plutonic rocks or in regionally metamorphosed rocks (Fig. 1.185). Deposits occur along the Tectonic lines such as Hidaka (Hokkaido), Kitakami (northeast Honshu), Tanakura (middle Honshu) and Itoigawa-Shizuoka (middle Honshu) Lines (Fig. 1.185). The deposits occurring in the sedimentary rocks (mainly black shale) are distributed mainly in three districts: Kitakami, Yamizo and Koma (Fig. 1.185). The few deposits in the metamorphic region are the Suwa, Kinkei, Amo and Hashidate (Fig. 1.185). Ages of mineralization in the Hidaka and Kitakami regions may be Cretaceous, considering the ages of associated granitic rocks. Isozaki and Maruyama (1990) thought that the Tanakura Tectonic Line moved at middle Miocene (ca. 15 Ma) related to the opening of Japan Sea. The time of movement of Itoigawa-Shizuoka Line is thought to be also middle Miocene. Ages of granitic rocks near the itoigawa-Shizuoka Line are middle Miocene (Shimizu, personal communication, 1998). These lines of evidence suggest that ages of mineralization along the Tanakura and Itoigawa Tectonic Lines are probably middle Miocene.
Miocene-Pliocene Hydrothermal Ore Deposits
251
AURIFEROUS VEIN Type 1
F7~J
60 Type 2
50
~
f°
after Nedachi (1974) after Yarnaoka (1981)
GOLD-SILVER VEIN E~ after Shikazono (1981) >, 40 O c-
$ u_ 30
20
10
0
0
20
40
60
80
100
Ag atomic % Figure 1.186. Frequency (number of analyses) histogram of Ag content (atomic %) of auriferous vein and gold-silver vein deposits in Japan. Frequency means numbers of analyses (Shikazono and Shimizu, 1987).
Main opaque minerals include native gold, electrum, pyrite, pyrrhotite, chalcopyrite, cubanite, sphalerite, arsenopyrite and tellurobismutite. The amounts of these sulfide minerals are poor, compared with those in epithermal A u - A g vein-type deposits. It is noteworthy that silver minerals are abundant in epithermal A u - A g vein-type deposits, whereas they are poor in gold-quartz veins. The Ag content of electrum is very low (Fig. 1.186) and FeS content of sphalerite is high (6-17 FeS mol%) (Fig. 1.187) (Shikazono and Shimizu, 1987). Combining these compositional data with homogenization temperatures of fluid inclusions, fs2 of ore fluids was estimated (Fig. 1.188). Estimated fs2 range is lower than that of epithermal A u - A g vein ore fluids. Shikazono and Shimizu (1987) calculated Z A u / I ; A g of ore fluids using fluid inclusion data (salinity, temperature), pH, mH2s and mK+. The procedure for estimating Z A u / Z A g by them is given below.
Chapter 1
252
100
•
AURIFEROUSVEIN
[--] GOLD-SILVERVEIN 80
O ¢-
60 t:r
1.1_ 401
20
--
0
2
4
6
8
10 12 14 16 18
FeS mole % of sphalerite Figure 1.187. Frequency (number of analyses) histogram of FeS content (mol %) of sphalerite from auriferous vein deposits in Japan (Shikazono and Shimizu, 1987).
^~,'-~*~//7..: . ~
~,
-10 d
-15 ,
~O~:5,,,,,.,~ ~
150
~' S,4~.,~
~'3P''f:££'O'a
200
250
300
350
Temp (°C)
Figure 1.I88. Typical sulfur activity and temperature ranges for Japanese auriferous vein (dotted) and goldsilver vein (hatched) deposits. Iso-FeS content curves for sphalerite were drawn based on the equation of Barton and Skinner (1979). py: pyrite, po: pyrrhotite (Shikazono and Shimizu, 1987).
It is essential to know the mode of transport o f Au and Ag in ore fluids to consider the factors which control the A g / A u ratio o f native gold and electrum. Many studies on Au and Ag complexes in ore fluids have been conducted and reviewed by several workers (Barnes and Czamanske, 1967; Barnes, 1979; Seward, 1981; Shenberger, 1986).
Miocene-Pliocene Hydrothermal Ore Deposits
253
According to these previous studies, the most dominant dissolved states of Au and Ag in ore fluids are considered to be bisulfide and chloride complexes, depending on the chemistry of ore fluid (salinity, pH, redox state, etc.). However, very few experimental studies of Au solubility due to chloride complex and Ag solubility due to bisulfide complexes under hydrothermal conditions of interest here have been conducted. Thus, it is difficult to evaluate the effects of these important species on the Ag/Au of native gold and electrum. Other Au and Ag complexes with tellurium, selenium, bismuth, antimony, and arsenic may be stable in ore fluids but are not taken into account here due to the lack of thermochemical data. Assuming that Au is transported dominantly as bisulfide or chloride complexes, the following reaction can be used to determine which species are dominant under near neutral conditions• AuC12 + 2H2S -----Au(HS) 2 + 2C1- + 2H +
(1-84)
The equilibrium relation for reaction (1-84) is expressed as 2 2 2 aauci 2/aau(HS) ~ = mauci 2/mAu(HS) ~ = (aci- all+)/(K1-84aH2S)
(1-85)
where a is activity, m is molality, and K1-84 is the equilibrium constant for equation (1-84)• It is assumed that the activity coefficient ratio, YauC12/yAu(US)2, is one. Several investigations to elucidate the geochemical environment of Japanese epithermal A u - A g vein-type deposits have bee,, ,'onducted (Shikazono, 1974a, 1977a, 1978b, 1985a,b). Based on these studies, the valu306of mauci-/mAu(HS) as a function of tem• perature may be calculated (Fig. 1.189), by taking the average values for a c v , all+, au2s, and temperature. The ratio of mAuci- /mAu(HS) increases with increasing temperature and acl (Fig • 1" 189), while an increase m pH causes a decrease in m I-A u t~. / m .~ U. [.r l.~.) ~. ratio " ~l 2 It is apparent in Fig. 1.190 that Au bisulfide species are more abundant than Au chloride species under the conditions common for ore fluids responsible for Japanese A u - A g veins. However, Au chloride species may dominate Au bisulfide species in ore fluids responsible for the gold-quartz (auriferous) vein deposits, as shown in Fig. 1.189. Assuming that AgC12 and Au(HS)~- are the predominant Ag and Au species, the following reaction may be used to derive the relationship between the Ag/Au ratio of native gold or electrum and temperature or other variables: •
•
.
.
2
.
2
2
(Au) + AgC12 + 2H2S = (Ag) + Au(HS)~ + 2C1- + 2H +
2
.
•
(1-86)
where (Au) and (Ag) are the Au and Ag components of native gold and electrum, respectively. From the equilibrium relation for equation (1-86) we obtain,
aAg/aAu ----(a22smagc12 KI-86)/mau(Hs)2a21 a2+
(1-87)
where K1-86 is the equilibrium constant for equation (1-86)• Equation (1-87) implies that the aAg/aAu of native gold and electrum is controlled by temperature, an.-, aN-S, pH, and m -A g t A~.2 / m A. .u t.t a.~.) 2. " ~1 2 The curves representing the relationship between temperature and aAg/aAu (aag, activity of Ag component in native gold or electrum; aAu, activity of Au component in native gold or electrum) are shown in Fig. 1.190. It is assumed that aK+/aN+ is controlled
Chapter 1
254
6 5
t~
3
iI::::: : ' ' ~ : : : :"::::::::' :": : :
z
£
fi
T-,
0
......
131 O --
-2
-3 -4 -5 ,
200
,
,
.
,
250
.
.
.
.
i
300
T(oC )
Figure 1.I89. The relationship between m .... /m . . . . . . . and temperature• Hatched and dotted areas represent • • , AUk.12 . AU[IJ~)2 the probable geochemical envlronment for typical Japanese gold-silver vein and auriferous vein deposits, respectively• A, mcl = 10, mK+ = 2, aH2S = l0 3, K-feldspar/K-mica/quartz equilibrium; B, reel = 1, mK+ = 0.2, aH2S = 10 -35, K-feldspar/K-mica/quartz equilibrium; C, mci = I, inK+ = 0•2, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium; D, mcl- = 0.2, inK+ = 0•04, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium; E, mcl = 0.2, inK+ = 0.04, aH2S = 10 -3, K-feldspar/K-mica/quartz equilibrium; F, mcl = 0.2, mK+ = 0.04, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium• Thermochemical data for the calculations were taken from Helgeson (1969), Seward (1973), Drummond (1981), and Henley et al. (I984)• (Shikazono and Shimizu, 1987).
by the K-feldspar-K-mica-quartz assemblage which commonly occurs in Au-Ag vein deposits in Japan (Shikazono, 1974a). The values of HzS activity in ore fluids responsible for A u - A g veins are assumed to be 10 - 2 - 1 0 - 3 , which are estimated from sulfide mineral assemblage and chemical compositions of minerals (e.g., Fe content of sphalerite; e.g., Shikazono, 1974a). The transport of Au in the ore fluids responsible for the gold-quartz veins probably takes place as Au chloride complexes rather than as Au bisulfide complexes (Fig. 1.190). This suggestion is based on the fact that the temperature of formation of gold-quartz vein deposits is high (probably 250-350°C) and the chloride concentration of the ore fluids is high (probably more than 1 molal and less than 10 molal) based on the fluid inclusion studies (e.g., Nedachi, 1974; Shikazono, unpublished). Based on the experimental data on the solubility of Au due to chloride complexes (Henley, 1984) Au is highly soluble at high temperatures and/or in fluids of high chloride concentration. In fluids where precious metal transport is dominated by AgCI 2 and AuC12, the following reaction may be written to consider the compositional variation of native gold or electrum: (Au) + AgCI~- = (Ag) -t- AuCI~-
(1-88)
Miocene-Pliocene Hydrothermal Ore Deposits 5
255
A
4 3
C
~2
%1 t~
~o,o
-1 -2 -3 200
250
300
Temp. (°C) Figure 1.190. The relationship between aAg/aAu of native gold and electrum and temperatures. A, aH2s = 10 2, mcl_ = 0.2, inK+ = 0.04, aAuC]£/aAu(HS)£= 102"°, K-feIdspar/K-mica/quartz equlibrium; B ' aH~S = 10 - 2 , m c i - = 0.2 , mK+ = 0.04 ' aAucI /aAu~HSa = 10 T M , C , aH-S = 10 -3 ~ mcl = 0.2 , z ~ /2 z mK+ = 0 . 0 4 , aauCl~-/ a Au(HS)~ = 10 2.0., D, a H 2 S = 10 -~ , me1 = 0 . 2 , mK+ = 0 . 0 4 , aAuCl~/aAu(HS)2 = 10 1.5., E, aAuCI~/aAu(HS)~= 1020; F, aAuci2/aAu(HS)~= 101"5. Calculations were m a d e a s s u m i n g K - f e l d s p a r / K m i c a / q u a r t z equilibrium for A, B, C, and D. T h e r m o c h e m i c a l data for the calculations were taken f r o m Helgeson (1969), Seward (1973, 1976), D r u m m o n d (1981), a n d Henley et al. (1984). (Shikazono a n d Shimizu, 1987).
From the equilibrium relation for equation (1-88) we obtain
aAg/aAu ~ ( a A g c 1 2 K1-88) /aAucI 2 (]-89) Curves E and F in Fig. 1.190 were drawn assuming that the value of N A g / E A u (NAg, total dissolved Ag concentration, NAu, total dissolved Au concentration) is approximately 102-1015 , which corresponds to the ratio for crustal rocks. These curves indicate that the aAg/aAu ratio of native gold or electrum does not change with temperature at a constant E A u / N A g ratio. It is noteworthy that the temperature dependency of the aAg/aAu ratio of native gold and electrum in AuC12 dominant ore fluids (Fig. l. 190). This difference may account for the range of variation and the Ag/Au ratios of native gold and electrum from gold-quartz vein and epithermal A u - A g vein deposits. However, it has to be noted that uncertainties of thermochemical data on gold chloride species might be large. As suggested by Henley et al. (1984) and Brown (1986), it is possible that Ag(HS)2 is dominant in low-salinity and near neutral ore fluids. Under this condition the Ag/Au ratios of electrum and native gold are controlled by the following reaction: (Au) + Ag(HS)~ = (Ag) + Au(HS)2
(1-90)
It is expected from the equilibrium ratios for equation (1-90) that the Ag/Au ratios of native gold and electrum are controlled by temperature and E A g / Z A u in ore fluids, although the equilibrium constant for equation (t-90) has not experimentally been determined. The E A u / E A g ratio in ore fluids responsible for epithermal A u - A g veins could be calculated from equation (1-87). The temperature of formation is taken from fluid
Chapter 1
256 N I
Fukushirr~ K.oganesawa
-~.~Oshima Tochigi Pref.~.'~'--. . ~ ~
Pref.
rune
om, o
Nokoz~mo...{/2/~."
Nosu~ ~ ' . " .... ~_-~z~"'o
• / ' . - Y.,
[: "("Ibaragi Pref. . .) H .) i--Tarozawa Tokakum~... - . . ~ t~etsukuji Otori_.===:/V~N..t---'/:~-'.-. . ." ~--'~Hmaga BATO,~J - .'X.) • xx[ . ,'..qh-.~Shiozawa ~.."-'~.;y /..~'_-_-_-_-_-_-_-.~Kuryu
a UTSUNOMIYA
Figure 1.I91. Locationof gold vein-typedeposits in the YamizoMountains. Dotted area represents the Yamizo Mountains except the Toriashi and Tsukuba blocks. Open circle: Au vein-typedeposits, solid circle: studied Au vein-typedeposits, square: city or town (Shimizu and Shikazono, 1987). inclusion studies (Enjoji and Takenouchi, 1976; Shikazono, 1985b) and the electrumsphalerite-argentite-pyrite assemblage (Shikazono, 1985d). The NaCI equivalent concentration of ore fluids is approximated from freezing data on inclusion fluids (Enjoji and Takenouchi, 1976), though the final melting temperature of fluid inclusion ice is also affected by CO2 concentration in epithermal ore fluids (Hedenquist and Henley, 1985). The pH values are estimated assuming the equilibrium among K-feldspar, K-mica, and quartz, which occur commonly in these deposits; this in turn allows a calculation of the activity of K +. The activity of HzS is estimated based on the equation showing the relation between the partial pressure of H2S gas and temperature for active geothermal waters (Giggenbach, 1980; Arn6rsson, 1985). Using the typical values of these variables and XAg = 0.5, which is a typical value for electrum from Au-Ag veins, E A u / E A g is calculated to be about 10 - I , which is similar to that of Broadlands geothermal water (Table 1.27). The above calculation is based on the assumption that AgCI~- is the predominant Ag species in ore fluids. However, Brown (1986) and Henley (1985) suggested that silver bisulfide complex (Ag(HS)2) could contribute significantly to the transportation of Ag in low-salinity geothermal waters (e.g., Broadlands, New Zealand). Thus, Ag bisulfide complex is also probable as dominant Ag species in ore fluids responsible for Au-Ag veins in Japan, though the salinity of ore fluids (0.1-0.3 molal) is generally higher than that of Broadlands geothermal water. If Au and Ag bisulfide complexes are the dominant Au and Ag species in Broadlands geothermal water and ore fluids responsible for Japanese gold silver veins, the Ag/Au ratio of ore fluids responsible for Japanese
Miocene-Pliocene Hydrothermal Ore Deposits
257
TABLE 1.27 EAu/NAg in ore fluidsestimated from chemical compositionof nativegold and electrum in active geothermal waters and in crustal rocks. XAg (X: atomic fraction) = 0.2 for auriferous veins and XAg = 0.5 for gold-silver veins are assumed. Assumed values include: (1) 200°C, aH2S 10-35, mcl 0.1, pH = 5.6; (2) 250°C, aHzS = 10 2, mcI =0.1, pH = 5.4; (3) 300°C, aH2s = 10-1, mcl- =0.i, pH = 5.4; (4) 300°C; (5) 300°C =
=
Au/Ag log (atomic ratio) (Au/Ag)
References
0.03~0.01 0.013~0.004
-1.6-2.0 -1.1-2.2
Wehdepohl(1978) Holland(1978)
0.i0 0.11 0.03
-0.99 - 1.00 - 1.6
Brown (1986) Henley et al. (1984) Koga (1957, 1961)
0.09
-1.0
(2)
0.09
- 1.0
(3)
0.08
- 1.1
Thermochemicaldata from Helgeson (1969), Seward (1973, 1976), Henley et aI. (1984) Thermochemical data from Helgeson (1969), Seward (1973, 1976), Henley et aI. (1984) Thermochemical data from Helgeson (1969), Seward (1973, 1976), Henley et al. (I984)
Auriferous vein (4)
0.007
--2.2
(5)
1.48 × 10 7
-6.8
(6)
0.007
-2.2
Crustal rocks (average) Seawater Active geothermal water
Broadlands (BR 22) Imperial Valley Beppu Ore fluids
Gold-silver vein (1)
Thermochemical data from Drummond(1981), Seward (I976) Thermochemical data from Helgeson (1969), Seward (1976) Thermochemical data from Drummond(198I), Seward (1976)
gold-silver veins may be similar to that of Broadlands geothermal water which is about 0.1 (Table 1.27). If AgCI~ and AuC12 are the predominant Ag and Au species, we can calculate E A u / E A g in ore fluids by using thermochemical data on gold chloride complexes by D r u m m o n d (1981), Helgeson (1969), and thermochemical mixing properties of A u - A g alloy by White et al. (1957)• The ratio of XAg/XAu (X: atomic fraction) for native gold or electrum from gold-quartz vein deposits is taken to be 0.25. The calculated value of m AuCl£/mAgCl2 based on thermochemical data on AuC12 by D r u m m o n d (1981) is about 10 -2. Using thermochemical data for Au chloride complexes compiled by Helgeson (1969), we obtain mAuci /mA_C1 which is very different from the A u / A g ratios of • 2 g 2 crustal rock and active geothermal waters (Broadlands, New Zealand; Imperial Valley Magmamax 1, California; Beppu, Japan), though the A u / A g values for the Imperial Valley Magmamax 1 and Beppu geothermal water do not reflect those in the deep fluid. ~348 of sulfide (pyrite) from this type of deposit varies widely, ranging from - 5 to +7%0 (Shikazono, unpublished), but average ~34S value is 0%0, suggesting igneous origin from deep part, but some from reduction of seawater sulfate in sediments. 313C values of
258
Chapter i
carbonates from these deposits are mostly from -5%0 to -7%0, suggesting igneous origin (Shikazono, unpublished).
1.6.3.2. Gold-quartz vein-type deposits in Yamizo Mountains, central Japan Shimizu and Shikazono (1987) studied gold-quartz vein-type deposits in Yamizo mountain, central Japan. Their study is summarized below. The Yamizo Mountains extend over the eastern part of Tochigi Prefecture, and the western part of Ibaragi Prefecture, central Japan, occupying an area ca. 20 km by 50 km. A lot of hypo/mesothermal (in the sense of Lindgren, 1933) gold vein-type deposits exist there (Fig. 1.191). The Yamizo Mountains are mostly occupied by Paleozoic-Mesozoic sedimentary rocks, mainly of Jurassic age (e.g., Sashida et al., 1982) with a small amount of intrusive granitoids of unknown ages. The Paleozoic-Mesozoic sedimentary rocks have been called the Yamizo Formation (Kanomata, 1961). It is composed chiefly of shale, sandstone, alternating beds of shale and sandstone, and a small amount of limestone and chert. The succession of the geologic units and geologic structure of the Yamizo Formation have been left pending due to complex structure such as upturned beds (Kasai, 1978) and submarine land sliding (Aono et al., 1985). Many small gold-quartz veins occur in the Yamizo mountains (Fig. I. 191). Some of them were already undertaken under the Satake clan in the Tokugawa era (more than 150 years ago). Most of them were productive during the 1930s to 1950s in several ten tons per month with a grade of ten to 50 grams per metric ton of gold and silver, respectively. Gold-quartz veins ranging in width from several centimeters to several ten centimeters are situated in shale and/or sandstone of the Yamizo Formation and rarely in granitoids. The strike and dip of the veins are various. For example, the veins of the Kuryu, Shiozawa, and Daigo deposits strike N40°W, N20-30°W, and N30°E, respectively, and dip 60°E, 30-50°W, and 20°W, respectively. The veins are composed mostly of quartz and a small amount of sulfide minerals (pyrite, pyrrhotite, arsenopyrite, chalcopyrite, sphalerite, and galena), carbonate minerals (calcite, dolomite) and gold, and include breccias of the host rocks with carbonaceous matters. Layering by carbonaceous matters has been occasionally observed in the veins. Banding structure, wall rock alteration and an evidence of boiling of fluids that are commonly observed in epithermal veins have not been usually found. Gold is present either as isolated grains in quartz or in association with sulfide minerals. In the case of the latter, gold is generally in direct contact with pyrite and/or galena. Gold is irregular in shape with 0.01 to l mm in size. Some of gold grains can often be observed by naked eyes. Analytical results on gold are shown in Table 1.28. Chemical compositions of gold are relatively homogeneous and low Ag contents. Frequency histograms for Ag atomic% of gold (Nag) from the Yamizo Mountains and other regions are summarized in Fig. 1.192. Values of NAg of gold from the Yamizo Mountains range from 5.9 to 18.4, and this is almost the same as those from the other regions (Fig. 1.192). However, NAg of hypo/mesothermal gold vein-type deposits in metamorphic rocks tends to be slightly lower than that in sedimentary rocks.
Miocene-Pliocene Hydrothermal Ore Deposits
259
TABLE 1.28 Representative chemical compositions of gold from the Yamizo Mountains, Japan (Shimizu and Shikazono, 1987) Ore deposit
Weight % Au
Atomic % Ag
Cu
Total
Au
Ag
Cu
Ag/Au
Kuryu
96.70 96.31 96.15 96.13 96.20 95.66 95.89 96.05 95.36 95.24
3.31 3.41 3.50 3.6I 3.77 3.79 3.90 3.97 4.12 4.26
0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00
100.01 99.72 99.65 99.74 99.97 99.46 99.79 100.02 99.48 99.50
94.11 93.93 93.78 93.58 93.33 93.22 93. l0 92.98 92.69 92.45
5.89 6.07 6.22 6.42 6.67 6.74 6.90 7.02 7.31 7.55
0.00 0.00 0.00 0.00 0.00 0.04 0.00 0.00 0.00 0.00
0.063 0.065 0.067 0.069 0.071 0.072 0.074 0.076 0.079 0.082
Shiozawa
92.83 92.10 92.24 92.15 92.22 92.62 92.20 91.96 91.55 91.64
7.00 7.06 7.11 7.16 7.24 7.37 7.37 7.45 7.48 7.55
0.05 0.05 0.02 0.04 0.06 0.03. 0.06 0.06 0.04 0.00
99.88 99.21 99.37 99.35 99.52 100.02 99.63 99.47 99.07 99.19
87.76 87.60 87.61 87.47 87.32 87.24 87.12 85.98 86.92 86.92
12.09 12.25 12.33 12.42 12.52 12.67 12.71 12.86 12.96 13.08
0.15 0.15 0.06 0.11 0.17 0.09 0.17 0.17 0.11 0.00
0.138 0.140 0.141 0.142 0.143 0.145 0.146 0.150 0.149 0.150
Daigo
92.32 91.68 91.65 91.77 91.54 91.64 91.27 91.41 91.33 91.16
7.54 7.5I 7.69 7.78 7.83 7.93 8.04 8.11 8.18 8.37
0.02 0.01 0.02 0.02 0.02 0.03 0.02 0.04 0.00 0.04
99.88 99.20 99.36 99.57 99.39 99.60 99.33 99.56 99.51 99.57
86.97 86.96 86.66 86.55 86.44 86.28 86.10 85.96 85.95 85.54
12.97 13.00 13.28 13.40 13.50 13.63 13.85 13.93 14.05 14.35
0.06 0.04 0.06 0.06 0.06 0.09 0.06 0.11 0.00 0.11
0.149 0.149 0.153 0.155 0.156 0.158 0.161 0.162 0.163 0.168
Saigane
89.73 89.40 89.56 88.69 89.19 89.12 88.42 88.99 88.30 88.34
9.68 9.71 10.19 10.28 10.46 10.54 10.66 10.78 10.87 10.91
0.07 0.06 0.06 0.06 0.08 0.08 0.07 0.08 0.08 0.09
99.48 99.17 99.81 99.03 99.73 99.74 99.15 99.25 99.25 99.40
83.38 83.31 82.67 82.39 83.17 82.05 81.79 81.46 81.46 81.31
16.42 16.52 17.17 17.44 17.59 17.72 18.01 I8.30 18.30 18.44
0.20 0.17 0.16 0.16 0.24 0.24 0.20 0.24 0.24 0.25
0.197 0.198 0.208 0.212 0.214 0.216 0.220 0.225 0.225 0.225
Chapter 1
260 N=144
In Metamorphic R. Suwa, Kinkef, Amo, Hashidate
50 4,0 30 20 10
El N: 1[
70
Sh
6O 5O
In Sedimentary R. Yamizo Mtns, Kuryu, Daigo Shiozawa, Saigane
D
o 40 o-
~-
30
u_ 2O
10 ~
~Sa
S. Koma Region Ho Gohaku, Koei
[
N=1'57j~ 30. A~erlgel)
' Kitakami Mtns. Oya, Shishiori,
20
Hayachine, Kohoku Nedachi(1974) Yamaoka(1981)
10
0
10 20
50
100 NAg
Figure 1.192. Frequency (no. of analyses) histograms for Ag content (in atomic %) of gold from "hypo/ mesothermal" gold vein-typedeposits in Japan. K: Kuryu deposit, Sh: Shiozawa deposit, D: Daigo deposit, Sa: Saigane deposit. Data are from: the Yamizo Mountains:Shikazonoand Shimizu (1988a); the South Korea Region: Shikazono and Shimizu (1987); the Kitakami Mountains: Nedachi (1974); Yamaoka (1981), Abe (1981) and Shikazono and Shimizu (1987); metamorphicregions: Shikazono and Shimizu (t987).
In the Yamizo Mountains, the correlation between NAg and a kind of the host rocks is found: NAg of gold from gold-quartz veins in shale is lower than that in sandstone. Gold negatively correlates with 834S of the associated sulfide minerals from the veins, and that 334S of sulfide minerals in sandstone is lower than that in alternating beds of shale and sandstone, probably than that in shale.
Miocene-Pliocene Hydrothermal Ore Deposits
261
Homogenization temperatures of fluid inclusions mainly in quartz closely associated with gold and sulfide minerals in hypo/mesothermal gold vein-type deposits in Japan are in a range from 250 ° to 350°C (Shikazono and Shimizu, 1987), higher than those in epithermal gold-silver vein-type deposits in Japan. Iron contents of sphalerite and thermochemical data by Barton and Toulmin (1966) and Barton and Skinner (1979), possible sulfur fugacity (fs2) of the ore fulids can be estimated: log fs2 = - 1 5 at 250°C to - 3 at 350°C. This means that fs2 for hypo/mesothermal gold vein-type deposits is lower than that for epithermal gold-silver vein-type deposits in Japan. Factors in controlling chemical compositions of gold in equilibrium with the ore fluids are temperature, pH, concentration of aqueous H2S and C1- in the ore fluids, concentration ratio of Au and Ag species in the ore fluids, activity coefficient of Au and Ag components in gold, and so on (Shikazono, 1981). In the Yamizo Mountains, as a result, Ag/Au ratios of gold are correlated with a kind of the host rocks and sulfur isotopic compositions of the deposits. This correlation could be used to interpret Ag/Au ratios of gold. There are two possibilities here to explain this correlation. One is that isotopically heavy sulfide sulfur derived from seawater sulfate was fixed in shale because reducing agency of shale with carbonaceous matters is thought to be stronger than that of sandstone. The ore fluids extracted this sulfur. Gold of low NAB precipitated in shale like the Kuryu deposit under more reducing environment than in sandstone like the Saigane deposit. Shikazono and Shimizu (1987) concluded that Ag contents of gold precipitated from low-salinity fluids is higher than that prediction and the relationship between Nag of gold and salinity of fluid inclusions estimated from freezing temperature data. Therefore, another interpretation is that Nag of gold from shale-hosted deposits is lower than that from sandstone-hosted deposits, because shale is expected to be richer in C1- mainly due to adsorption by clay minerals included in shale than sandstone. It is concluded that, in either case, in situ interaction of the ore fluids with the host rocks controls the chemical compositions of gold in the Yamizo Mountains.
1.6.4. Hot spring-type gold deposits Hot spring-type gold deposits (Nansatsu-type by Urashima et al. (1981, 1987), high sulfidation-type by Hedenquist (1987), epithermal Au disseminated-type) are distributed in the Nansatsu district of southern Kyushu (Fig. 1.193). The deposits (Kasuga, Akeshi, Iwato) were formed at Pliocene age (5.5-3.7 m.y.) in the calc-alkaline volcanic rocks of nearly same age (Togashi and Shibata, 1984). The deposits, which are similar to Nansatsu-type deposits, occur in Southwest Hokkaido (Date, Hakurhu). The deposits are characterized by conspicuous alteration zoning from the centre (orebody) to margin (Tokunaga, 1955; Doi, 1972; Urashima et al., 1981, 1987). They are siliceous zone, alunite zone, kaolinite zone, sericite zone and montmorillonite zone. In the siliceous body, electrum and Cu minerals (enargite, luzonite, covelline), and native sulfur occur. The Ag content of electrum is lower (0.0-5.3 wt%) than that from epithermal A u - A g vein-type deposits (Fig. 1.194) (Shikazono and Shimizu, 1987). Low Ag content of electrum and sulfide mineral assemblage (enargite, native sulfur, covellite, pyrite) indicate high i s 2 condition (Fig. 1.194).
Chapter I
262
J~ f J
~ ~
\ o
~._~v-
~
[ [
- ~
~
• Hot springtype
~ EpithermalAu-Agvein-type ~ Active volcano
o
20
I,
~,
4o
I
/
o o~uo~i
)
/
o,.o..=o
(
Kagoshima
/
j
Mt.Kaimondake
31N
I
~
l 131E
Figure 1.193. Locality map of the hot spring-type deposits in Kyushu. 4O
=~ 3o
g
2O
10
0
0
20
40
60
80
1oo NAg
Figure 1.194. Frequency histogram for the Ag content of eIectrum from epithermal Au disseminated-type deposits in Japan (Iwato and Kasuga). (Shikazono and Shimizu, 1987).
Miocene-Pliocene Hydrothermal Ore Deposits
263
/
Steam-heated alunite fluid / / / I .i~<-~/
•
K ~
I 0 Alunite
•
0 Clays
SMOW
.~'V Alunitefluid '@ ~
{ 1._.: ~s~ ~
volcanic fluids
O
-50 O-shift
a
Clayfluid 0 ~ 0 0
-100
.2
I
I
t
I
Residual silica • I
(
I
H
•
silica
Kasuga
and
0
O Iwato
Deepveins at Kasuga
(~
Peripheralveins Kago veins
~
vein
quartz
Other localities Calculated fluids
I
-10
•
0 Residual
Quartz veins in residualsilica
I
0
8180
+10 %o
I
+20
Figure 1.I95. Diagram showing the results of isotopic analysis of residual silica (now recrystaltized to quartz) and quartz crystals or veins (3180 only, plotted on the bottom), and clays and alunite (~D vs. 3180, plotted on the top). The calculated composition of the fluids in equilibrium with these minerals, plus the compositions of local meteoric water (small open square) and high-temperature volcanic fluids (large open square) are also shown (Matsuo et al., 1974; Matsubaya et al., 1975). A least-squares fit line (dashed) through the hypogene alunite fluids and the high-temperature volcanic fluid end member extends to an isotopic composition similar to present local meteoric water and identical in gD values to the clay fluid. Thus the isotopic composition of alunites record mixing of magmatic and meteoric waters, whereas the marginal clays (and possibly the vein quartz and residual silica quartz) indicate that meteoric water dominated on the margins of the quartz bodies, variably shifted in oxygen isotope composition due to exchange with wall rocks. Some of the deep fluids responsible for vein formation are identical in isotopic composition to the 180-enriched fluid that equilibrated with the residual silica quartz, whereas other samples indicate a greater proportion of unshifted meteoric water. One vein alunite from Iwato has an isotopic composition which indicates that it formed in a steam-heated environment after evaporation of local ground water (fluid composition calculated for a steam-heated temperature range of 100-180°C). K and I denote Kasuga and Iwato, respectively. (Hedenquist et al., 1994).
Chapter 1
264
Hedenquist et al. (1994) indicated that homogenization temperature of postmineralization of the Akeshi deposit and salinity of ca. 1 wt% NaCl eq. and the presence at times of a two-phase fluid in the centre of ore zones indicating boiling of ore fluids. 834S of sulfides (pyrite, enargite, luzonite) and sulfate(alunite) are -5%0 to 0%0 and +24%0 to +35%0, respectively (Shikazono, 1987b; Hedenquist et al., 1994), suggesting that sulfide sulfur came both from igneous and sedimentary sources and sulfate sulfur is generated by the hydrolysis reaction of 4 S O 2 -k- 4H20 ~ H2S + 3 H 2 8 0 4 . SO2 gas may come from volcanic gas. 3D and 8180 of ore fluids were estimated from these data on quartz, clay minerals and alunite (Fig. 1.195) (Hedenquist et al., 1994), indicating higher 8D and 8180 values compared with epithermal Au-Ag vein ore fluids. These isotopic data (834S, 8D, 8180), fluid inclusion data and alteration studies suggest that (1) the gas-rich (HCl, SO2, H2S) vapor phase separated from magma ascent to the surface and condense into meteoric water, (2) this acid chloride-sulfate fluid leached elements from the host rocks, forming a permeable zone of mixing, (3) then, the dense metal-rich liquids ascent into the leached zone and precipitate Cu-minerals and Au (Hedenquist et al., 1994). The fo2-PH conditions for this type of deposits and other types of epithermal deposits and hydrothermal alteration are shown in Fig. 1.196 (Shikazono and Aoki, 1981). It is worth noting that the f o 2 - P H ranges for the epithermal Au deposits (hot spring-type, Te-type, and Se-type) lie between the points A (high fo2, and low pH) and B (low f o e
0
Log fo 2 'o~ vm, e High sulfidation type
'r ,
I
I
/ll
~AIIKal , / 1 , Y"
-30 •
."--. ,,
l./I
Low sulfidation type
/I /'
/ s
s,
-
Hm Mt -35
Bn+Py Cp
I . ~ t - " ",. I "3-. I
D,,
~
Po
--I---I
~40
i
I
Kal 1
I
I
2
3
4
I
5
2~
'
4
1
1
e4~CO
I I
~
-r ', -r Kf ' 1
I
Se I
~.
' ""!"';"c'--
I 1
1
I
I~
|
I
6
7
8
9
pH
Figure 1.196. fo2 pH ranges for hot-spring-type deposits and low sulfidation-type deposits. Temperature = 250°C, £S = 0.01 mol/kg H20, ionic strength = I. Ka: kaolinite, AI: alunite, SI: liquid sulfur, Kf: K-feldspar, Hm: hematite, Mt: magnetite, Py: pyrite, Po: pyrrhotite, Bn: bornite, Cp: chalcopyrite.
Miocene-Pliocene Hydrothermal Ore Deposits Assumed direction of the Pacific Plate Deviation form the north 90 ° 70 ° 50o 30 °
Ma 0
EpithermaI system
Q.R intrusion
~
O
12 -
f
N
3.
- Tokiwamatsu H.S.S. - Kobui H.S.S.
.
4 5-
265
-O
Q.R
Hokko-Minami LS.S.
- - Yamato L.S.S.
6--Mitsumoriyama
7 -
Q.R
H.S.S.
8N 910-
H.S.S.: High sutfidation s y s t e m L.S.S.: L o w sulfidation s y s t e m Q.R: Q u a r t z p o r p h y r y N: n o r m a l s u b d u c t i o n , O: o b l i q u e s u b d u c t i o n Figure 1.197. Subduction mode of the Pacific plate beneath the Northeast Japan arc, and style of the epithermal systems in the Kameda peninsula, Hokkaido. Normal, oblique and highly oblique subduction are tentatively defined by the angle (c0 between the assumed subduction direction of the Pacific plate and the trend of the Northeast Japan arc. Normal subduction (60 ° < c~ = 90°), oblique subduction (30 ° < c~ = 60 °) and highly oblique subduction (0° < c~ = 30°). Highly oblique subduction has not occurred along the Northeast Japan arc since 11 Ma. (Watanabe et al., I996).
and high pH) in Fig. 1.196. These difference in physicochemical nature of fluids can be explained by space-dependent evolution of a hydrothermal system relative to the centre of volcanic activity (Shikazono and Aoki, 1981; Shikazono, 1985a; Shikazono et al., 1990; Sillitoe, 1991) or a time-dependent evolution of a hydrothermal system (early-stage; high sulfidation-type, late-stage; low sulfidation-type) (Aoki et al., 1993). For example, Shikazono (1985a) has shown that advanced argillic alteration (Ugusu silica deposit in west Izu Penninsula, central Honshu) occurs at the centre and
Chapter 1
266
stratigraphically higher position in the mine district, while epithermal Au-Ag veintype deposits at the margin and lower position (Fig. 1.76). Shikazono et al. (1990) compared the characteristic features of Te-type with Se-type and suggested that the Te-type formed at the position closer to volcanic centre in a hydrothermal system than Se-type (Fig. 1.123). It is probable that the difference in fo2 and pH in several types of epithermal Au deposits is due to the different ratio of meteoric water component (low fo2 and high pH) to magmatic water (or igneous water) component (high fo2, low pH), related to the position from volcanic centre. Watanabe et al. (1996) and Watanabe and Ohta (1999) pointed out that the high sulfidation (advanced argillic alteration) and low sulfidation (sericitic and propylitic alteration) system in Southwest Hokkaido (Minamikayabe area, Jozankei-Zenibako district) occurred during normal and oblique subduction of the Pacific plate beneath the Northeast Japan arc, respectively (Fig. 1.197). These studies suggest that the difference in two styles of epithermal gold mineralization (epithermal vein (low sulfidation-type) and hot spring-type (high sulfidation-type)) is caused by the subduction mode of the plate and not by the temporal and spatial differences in a given hydrothermal system. Figure 1.198 shows grade-tonnage relationship for low sulfidation-type (Se-type) and high sulfidation-type (Te-type and Akeshi-type) in Japan and deposits in Circum Pacific region summarized by Hedenquist et al. (1996). This clearly indicates that grade and tonnages for Akeshi-type and Te-type in Japan are low and small, compared with low sulfidation-type in Japan. It is obvious that low sulfidation-type is dominant Au-Ag deposits in Japan. The plot of K20 versus SiO2 for volcanic rocks thought to be genetically related to Au mineralization indicates that high sulfidation deposits (hot spring-type deposits appear
ll-t
100
I
Illllll
I Xl~
"'.z°o, -. 7 "za
"
-.e %
""-
]0
o", ",
eO OO
lllllll
I
I
~
I
[[IFF[
I
I
F lIIII
• Low sulfidation • High sulfidation
%0
,.L.,
. ' ~ "" ~
•
•
""
" "•-o RO • • "~'oEmMc oIB~'" "PV La "e. • • p p
] 0.l
[
"" f.ooa "- z Hi "'.
1
10
,e. A 100
RM ". 1000
Tonnage, Mt Figure 1.198. Grade-tonnage plot for epithermal Au deposits. LS: Ba, Baguio (Philippines); CC, Cripple Creek; Mc, McLaughlin and RM, Round Mountain (USA); Era, Emperor (Fiji); Hi, Hishikari (Japan); Ke, Kelian (Indonesia); La, Ladolam and Po, Porgera Zone VII (Papua New Guinea). HS: Ch, Chinkuashih (Taiwan); EI, EI Indio (Chile); Go, Gotdfieid and PP, Paradise Peak (USA); Le, Lepanto (Philippines); PV, Pueblo Viejo (DominicanRepublic); Ro, Rodalquilar (Spain) (Hedenquist et al., 1996).
Miocene-Pliocene Hydrothermal Ore Deposits 8
i
J
i
267
i
i
t
[
• LS(calc-alkalic) • LS(shoshonitic-alkalic) • HS(caIc-alkalic)
6
jO 4~
4
4 ~
•~
• ~0
•
._
•
L-A
•
•
•
• High K
~
Medium-K
¢, •
I
O0
Low-K I
50
I
l
60 SiO2,wt%
I
I
70
I
80
Figure 1.199. Plot of K20 versus SiO2 for volcanic rocks thought to be genetically related to Au mineralization. High sulfidation (HS) deposits appear to be associated with a narrow range of igneous rock composition, dominated by dacite and andesite. Low sulfidation (LS) deposits are largely associated with andesite and rhyolite, but major deposits such as Ladolam, Porgera and Cripple Creek are associated with shoshonitic and alkalic rocks (Hedenquist et al., 1996).
to b e a s s o c i a t e d w i t h a n a r r o w r a n g e o f i g n e o u s r o c k c o m p o s i t i o n , d o m i n a t e d b y d a c i t e a n d a n d e s i t e , w h i l e l o w s u l f i d a t i o n d e p o s i t s are w i t h a n d e s i t e a n d r h y o l i t e , b u t m a j o r d e p o s i t s (e.g., L a d o l d m , P o r g e r a a n d C r i p p l e C r e c k ) are a s s o c i a t e d w i t h s h o s h o n i t i c a n d alkalic r o c k s ( H e d e n q u i s t et al., 1996) (Fig. 1.199).
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294
Chapter 1
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1.1. General overview of metallogeny and tectonics in the Japanese Islands Before mentioning the characteristics of Kuroko and epithermal vein-type deposits in Japan, it is worthwhile to briefly describe the metallogeny, geology, geophysics, and tectonic situations of the Japanese Islands. Japan is situated along the boundary between Eurasia and the Pacific plate (Fig. 1.1). The metallogeny, geology, geophysics and geochemistry of Japan are largely controlled by this tectonic setting. The great variety of mineral deposits of Japan reflects the complex geotectonic environments. An intimate relationship exists between igneous and hydrothermal activity, which in turn reflects the plate tectonic history of Japan. Many Japanese ore deposits have produced many different metals, and they contain almost all common and useful minerals, although many deposits are small in size. Important metallic ore deposits include Besshi (Kieslager)-type (strata-bound cupriferous pyritic deposits), strata-bound Mn-Fe-type, skarn-type, Kuroko-type and vein-type. Dominant non-metallic deposits are limestone, clay, native sulfur, zeolite, silica and gypsum deposits. The deposits are divisible into three groups, based on their ages of formation: Carboniferous-Jurassic, Cretaceous-Paleogene and Tertiary-present. Carboniferous-Jurassic deposits, closely associated with submarine volcanic rocks, are of two kinds: Besshi (Kieslager)-type, and strata-bound Mn-Fe-type deposits. Besshi-type deposits are cupriferous pyritic deposits and occur mainly in metamorphic terranes (Sanbagawa, Sangun, Abukuma and Hidaka) and in some other areas (Chichibu and Shimanto; Fig. 1.2). The geological and geochemical similarities of these deposits and modern midoceanic ridge deposits (e.g., Juan de Fuca ridge, Guaymas) suggest a similar origin. For instance, the sulfur isotopic compositions of both types of deposits are equal to or higher than mantle values (generally + 1%0 to +4%0), suggesting mantle origin, perhaps modified by seawater-basalt interactions. In the Chichibu Zone, the intimate association of abundant strata-bound Mn-Fe deposits, limestone-dolomite and silica (chert) with basic volcanic rocks suggests an ocean-ridge hydrothermal origin. Jurassic-Cretaceous Besshi-type and Mn-Fe strata-bound deposits are present in Hidaka, Hokkaido (Fig. 1.2). Geochemical data and geological evidence all point to a midoceanic ridge environment of ore formation. 334S values of Shimokawa Besshi-type
Chapter 1 North American Plate
Eurasian
. ,o
Plate
.::.,y;
Sea of J a p a n -'~
Pacific Plate
•"::'. -**oh
\ Basin
~I
• :.:.
Philippine Sea Plate ' ~--..~-"~ I
TROUGH
I
/ /
PARECE VELA BASIN
/ Figure 1.1. Outline index map of the Japanese subduction zones. Thick lines with teeth are converging plate boundaries. Arrows indicate relative plate motions. Abbreviations: su, Suruga trough; sa, Sagami trough; sf, South Fossa Magna triple junction; och, Off Central Honshu triple junction; ISTL, Itolgawa-Shizuoka Tectonic Line; KSM, Kashima VLBI station (Uyeda, 1991).
deposits range mostly from +7%o to +10%o (Miyake and Sasaki, 1980), suggesting a contribution of seawater sulfate in addition to mantle source sulfur. In C r e t a c e o u s - P a l e o g e n e time many skarn-type and vein-type deposits formed associated with intense granitic activity. Granitic rocks are divisible into magnetite-series and ilmenite-series (Ishihara, 1977). Magnetite-series granitoids are present in North Honshu (Kitakami) and in the inner zone o f Southwest Honshu (San-in) Belt and ilmenite-series granitoids in the outer zone of Southwest Honshu (San-yo) belt. Metals associated with these two types of granitic rocks are distinct: Mo, Cu, Pb, Zn, Au and A g with the magnetite-series; Sn, W and rare earth with the ilmenite-series. Isotopic (Sr, S and O) data suggest that the ilmenite-series granitic m a g m a was influenced by
~ 1
Cretaceous-Paleogeneaccretionarycomplex
~
2 Cretaceoushigh P/Tme|amorphicbelt
~
3 JurassicaccreationaPjcomplex
~
4 Paleozoicaccretionarycomplex
~ 5
Continentalblock
@Cu(Bessh!subtype) ,,
c- ( . i . . o h i ~.,k,,ype)
•
~
M.
%
-/.
•
,/i~ ~ (~ ~
Sanbagawabelt
0
"~
~
S.Kitakami
HONSHU
0
100
200km
Figure h2, Distribution of the stratiform Cu-su|fide and chert-hosted Mn deposits in Japan (Sato and Ka_~e, 1996). MTL: Median Tectonic Line; TTL: Tanakura Teetonic Line; ISTL: Itoigawa-Shizuoka Tectonic Line; BTL: Butsuzo Tectonic Line.
4
Chapter 1
contamination of carbon-bearing sediments, whereas the magnetite-series granitic magma ascended from the lower crust without a significant interaction with carbonaceous matter. In the Kitakami district (North Honshu) gold vein-type (mesothermal-type) deposits and Cu-Fe skarn-type deposits occur associated with magnetite-series granitic rocks. Ore deposits associated with volcanic rocks generally exhibit polymetallic (Cu, Pb, Zn, Sn, W, Au, Ag, Mo, Bi, Sb, As and In) mineralization. Sulfur isotopic values of sulfides from these deposits are close to 0%o, suggesting a deep-seated origin of the sulfide sulfur. Clay deposits (pyrophyllite, sericite and kaolinite) are associated with both felsic volcanic rocks and ilmenite-series granitic rocks of late Cretaceous age in the San-yo Belt. Tertiary-Recent mineralization in Northeast Japan includes both epithermal veintype and Kuroko deposits (Fig. 1.3). Kuroko deposits occur only in the Green tuff region, whereas vein-type deposits occur both in the Green tuff region and in subaerial Tertiary-Quaternary volcanic region (Non-Green tuff region). K-Ar ages of formation of Kuroko and vein-type are middle Miocene (13-t- 1 Ma) and Plio-Pleistocene (3 4-2 Ma), respectively. Precious metal vein-type deposits include enrichments in Au, Ag, Hg, Te, Se, Sb, As, S, and Bi. Base metal vein-type deposits contain Pb, Zn, Mn, Ag and Cu, whereas Kuroko deposits are enriched in Cu, Pb, Au, Ag and Ba. In Southwest Japan, two styles of vein-type mineralization (Hg and Sb) formed from middle Miocene to the present. Many Hg and Sb deposits are present along the Median Tectonic Line, associated with the Setouchi andesites and ilmenite-series granitic rocks (Fig. 1.3). These different sites of hydrothermal and ore-forming activity may have resulted from the mode of subduction of the Pacific Plate. Mariana-type subduction (characterized by a steep angle of subduction and back-arc basin formation; Uyeda and Kanamori, 1979) during middle Miocene caused WNW-ESE extension, submarine hydrothermal activity, thick accumulation of bimodal (basaltic and dacitic) volcanic activity (Green tuff) and Kuroko-type formation (Shikazono and Shimizu, 1993). Plio-Pleistocene Chilean-type subduction (shallow-dipping subduction zone, E - W compression; Uyeda and Kanamori, 1979) and oblique subduction of the Pacific Plate beneath the North American Plate led to uplift and expansion of land area, subaerial hydrothermal activity accompanied by meteoric water circulation, subaerial andesitic volcanic activity and formation of vein-type deposits. Figure 1.4 shows the quantities of metals produced f¥om the metallic deposits in Northeast and Southwest Japan. This figure demonstrates that a large quantity of base metals (Cu, Pb and Zn) and precious metals (Au and Ag) was concentrated in the deposits of Northeast Japan, while they are much less abundant in the deposits of Southwest Japan. Subduction of the westward-advancing Pacific Plate under Northeast Japan was active during the Miocene, while in Southwest Japan, subduction along the Nankai Trough began more recently (5-10 Ma) and the Miocene ocean plate was inactive. These different plate motions may cause differences in quantities of sulfide ores and kinds of metals concentrated in the deposits. Lead and sulfur isotope data suggest that during subduction, pelagic sediments and altered basalt were incorporated into the magma in the island-arc trench systems of Northeast Japan and that lead and sulfur in vein-type and
Miocene-Pliocene Hydrothermal Ore Deposits
5
8, P Zn,S,~, ( £ A~) <~ Nn H£ Sb gure~ @NStS Bed~ ma~sanesee~s~s
Figure 1.3. Distribution of mineral deposits and tectonic provinces of the Neogene in Japan (Tatsumi, 1970). I, Zone of Miocene volcanism (Green Tuff region; mainly submarine); II, Zone of Miocene volcanism in the Ryukyu Arc (mainly subaeriaI); III, Zone of Neogene volcanism along the Median Tectonic Line (mainly subaerial); IV, Zone of Late Neogene folding, mainly in the Green Tuff region; V, Zone of Neogene intrusives and extrusives in the Shimanto terrane. Kuroko deposits in Northeast Japan originated from these materials. Antimony, mercury and sulfur in the H g - S b deposits in Southwest Honshu may have been derived from the shallow level of the crust under the Shimanto Group. Large epithermal gold vein-type deposits occur at major a r c - a r c junctions (Figs. 1.5 and 1.6); specifically, Chishima (Kurile)-Northeast Honshu, Northeast H o n s h u I z u - B o n i n and Southwest H o n s h u - R y u k y u . This m a y result from hydrothermal activities and mineralizations caused by intense volcanism at the a r c - a r c junctions. Hydrothermal c l a y - s i l i c a deposits (kaolinite, halloysite, sericite, montmorillonite and silica) and zeolite deposits occur in Tertiary-Quaternary volcanic regions. These deposits are distributed in areas o f epithermal gold mineralization.
Chapter 1 Miocene-Pliocene
Quaternary
Isla2d HE
i
Valcanic ' ,I
Japan
SW dapa.~n.n 0
DRI 30000
1
20000
40000
Kuroko-type deposit
6000
U/ZA
2000
Vein-type deposits
20000 100000 1 2 3
NN
Sulfur deposits
Figure 1.4. Comparison of quantities of ore deposits formed in late Cenozoic in NE and SW Japan. Weight per kilometer length of island arc (lshihara, 1978).
Quaternary sulfur deposits are distributed along the present volcanic front. Intersections of transverse faults proposed by Carr et al. (1973) and the present volcanic front coincide with the locations of clusters of the sulfur deposits (Nishiwaki and Yasui, 1974). Recently, it was found that mineralization is taking place in and around the Japanese Islands: sulfide-sulfate chimneys were discovered at back-arc depressions of the Ryukyu Arc Okinawa (Trough) and Izu-Bonin Arc (Smith Rift). The geologic settings are similar to those of the Miocene Kuroko deposits. The Ryukyu Arc belongs to the Mariana-type because its back-arc region is under extensional stress and the Okinawa Trough probably is a nascent back-arc spreading basin (Uyeda, 1991). The Izu-Bonin Arc may also be a Mariana-type, at present. However, it is likely that the arc was a Chilean type because intense epithermal gold mineralization took place at 1-3 Ma in the Izu Peninsula. Gold-rich silica precipitates at the Osorezan volcano, which is located in the most northern part of Honshu, have features very similar to epithermal Te-bearing gold vein-type deposits of the Plio-Pleistocene.
1.2. General overview and classification of hydrothermai ore deposits of Neogene age Main hydrothermal ore deposit types of Neogene age that formed in and around the Japanese Islands are Kuroko deposits and epithermal vein-type deposits. This classification is based on the form of the deposits. Kuroko deposits are strata-bound and massive in form (Fig. 1.7) and syngenetically formed on the seafloor and/or sub-seafloor environment. Vein-type deposits are fissurefilling and epigenetically formed (Fig. 1.8). Elemental association can be used to sub-classify these deposits. Major metal elements produced from Kuroko deposits are Cu, Pb, Zn, Ba, Ca, Fe, Au, and Ag. Average ore grade and tonnage are summarized in Table 1.1. Horikoshi and Shikazono (1978) classified Kuroko deposits into three sub-types: C sub-type (composite ore type),
Miocene-Pliocene Hydrothermal Ore Deposits
Figure 1.5. Three island arc junctions in the Japanese Islands (Kubota, 1994).
Y sub-type (yellow ore type), and B sub-type (black ore type), according to Cu, Pb and Zn ratios (Fig. 1.9). However, the variation in the ratio is not wide, compared with epithermal vein-type deposits. Therefore, characteristic differences in each sub-type of Kuroko deposits are not discussed here. Major epithermal vein-type deposits in Japan are base-metal type and preciousmetal type which are classified based on the ratios of base metals and Au and Ag which have been produced during the past (Table 1.2). Base-metal vein-type deposits may be divided into Pb-Zn-type and Cu-type (Otsu and Harada, 1963). However, this sub-classification is not considered here for simplicity of discussion.
T A B L E 1.1 Size and composition of ore deposits in the Hokuroku basin (Tanimura et al., 1983) Deposit name
Discovery (year)
Size (max. length x width x thickness, m)
Average ore grade Cu (%)
Zn (%)
Ph
Fe
s
(%)
(%)
(%)
1.13 2.28 1.32 2.60 I. 15 1.50 3.02 1.00 2.32 2.82 2.70 2.39
1.4 1.3 10.0 7.8 1.3 3.0 3.0 6.5 6.5 2.4 1.5 3.6
0.4 0.2 1.8 1.0 0.2 0.5 0.5
12.8 20.4 5.2 13.4 8.9 15.8 20.0
1.2
%0
0.5 0.4 0.3 1.0
7.0 24.4 14.8 3.6
14.7 23.4 6.0 15,4 10.2 18.2 23.0 8.0 8.0 28.0 17.0 21.1
Shakanai ore group (Shakanai Mines Co., Ltd.) No. I 1962 3 0 0 x 1 5 0 x 12 No. 2 1963 ?x?x7 No. 3 1963 400x 120x 6 No. 4 1963 400 x 300 x 40 No. 5 1964 350 x 70 x 13 No. 7 1965 350 x 250 x 15 No. 8 1965 430 x 170 x 40 No. II 1967 4 0 0 x l l 0 x 10
2.3 2.3 1.1 1.7 1.9 1.3 0.7
14.6 12.3 10.0 2.9 3.4 3.2 1.0
3,2 7,6 6.2 0.7 1.0 0.9 0,2
12.2 8.8 10.4 19.1 14.8 22.6 28.7
Matsuki ore group (Mitsubishi Metal Corp.) Matsuki 1964 350 × 90 x 30 Takadate 1963 150 × 150 x 30 Takadate South 100 × 70 x 15
1.9 3,74
11.8 2.00
3.4 0.80
0.89
10.1
3.32
Hanaoka-Shakanai area Hanaoka ore group (Down Mining Co., Ltd.) Tsutsumizawa [912 150 × 80 x 1 l0 Doyashiki 1916 270 x 200 x 100 Kamiyama 1919 60 x 40 x 55 Nanatsudate 1929 100 x 40 × 50 Higashi Kannondo 1935 35 x 15 x 25 O y a m a 1 Nishi 1938 60 × 13 x 30 Kannondo 1939 40 x 35 × 45 lnarizawa 1940 40 x 20 x 15 Ochiaizawa 1941 50 x 35 x 20 Oishizawa 1941 50 x 40 x 20 Oyama 2 1941 ll0x 17x30 Malsnmine 1963 600 x 400 x 110
Ezuri-Fukazawa
All (ppm)
Ag (ppm)
m m
m
Tonnage (1,000 metric tons)
441 8,946 932 677 105 60 258 34 207 90 100 30,000
0.5
57
t4.0 10. l 12.0 22.0 17.0 26.0 33.0
2.0 1.7 0.8 0.3
270 260 410 25
15.7 21.2
18.0 24.38
0.6
55
660 1,200
4.1
4.7
1.2
180
3,000
540 360
3,600 430 1,000
2,800
area
Ezuri ore group (Dowa Mining Co., Ltd.) Ezuri 1975
T A B L E 1.1 (continued) Deposit name
~. Discovery (year)
Size (max. length × width × thickness, m)
Fukazawa ore group (Dowa Mining Co,, Ltd.) Tsunokakezawa 1 1973 500 x 300 x 5 Kanayama 1976 210 × 90 x 8 Manjyaku 1979 190 x 190 x 13
Average ore grade Ca
Zn
Ph
Fc
S
AH
(%)
(%)
(%)
(%)
(%)
(ppm)
1.13 1.6 1.0
15.4 19.0 10,1
3.3 5.8 1.5
4.4 7.9 3.6
5.1 9,1 4.1
0.6
Ag (ppm)
93
Tonnage (1,0130 metric tons)
e~ I
~z
3,000
,5 g-
K o s a k a area
Uwamuki ore gnmp (Dowa Mining Co., Ltd.) No. 1 I962 1 5 0 × 1 0 0 x 14 No. 2 1965 200 × 150 x 40 No. 4 1966 350 x 1 0 0 x 17
0.6 0.8 0.8
11.5 7.8 8.3
4.2 1.8 2.8
4.1 7.7 5,5
4.7 8.8 6.3
0,7 -
Uchinotai ore group (Dowa Mining Co., L t d ) West 1959 400 x 300 x 70 East 1960 400 x 300 x 40
2.8 2.0
4.0 4.4
1.1 1.5
17.2 13.6
19.8 15.6
0.8
Motoyama ore grcmp (Dowa Mining Co., Ltd.) Motoyarna 1861 300 x 700 x 50
2.2
4.5
0.8
20.6
23.7
Furutobe ore group (Mitsubishi Metal Corp.) Yunosawa Daikoknzawa 1959 250 x 1 0 0 x 15 Daikokuzawa West $960 t70 x 70 x 40 Daikokuzawa Easl 1960 100 x 60 × 20 M a g a r i y a z a w a East 1962 200 x 80 x 15 M a g a r i y a z a w a West 1962 150 x 50 x 15
1,9 1.1 2,8 1.5 1.5
4.3 1.0 6.2 0.9 0.9
0.9 0.1 1.4 0.2 0.2
17,4 15.3 24,4
20.0 17.6 28.0
A inai ore group Ytmosawa Suehiro Daikoku Benten Yokodawara Hagoromo
0,7 4.7 2,2 1.9 2.0 1.9
tr 8.3 5,1 3.1 2.9 10.3
lr 1.7 1.3 1.2 0.7 3.5
21.8 17.6 16.1 15.7 18.3 20.0
25.0 20,2 18.5 18.0 21.0 23.0
120
130
160
930 2,580
5,240 4,000
15,000
Furutobe-Ainai area
1942 1955 1956 1957 1960 1967
200 50 180 200 150 150
x x x x × x
150 x 50 40 x 30 80 × 50 60 x 20 80 x 40 80 x 10
1.3
51
3,800
t.0 0.5 0.4 9,1
260 140 130 620
220 2,000 1,500 210
Chapter 1
10
r",, 12.912.8 2.93-049( !," ~,~,.;" 4.9~2.7\ ~ ~P "'t'~',,T/'~..,, ~[~] 3[W-] 5~b] 13
50.t < lO-50t I ~lOt o 0.II I
5.8
o 501oo 200 300km 1 ' ~ 3 " J I
I
I
I
I
(4.6)
3.6
0.6 1.1,3.7
1.1 4 0~--1.0 514 , ~1.5~1.8 ~-3.7 -4.3 Figure 1.6. Distributionand temporal and spatial relationshipof late Cenozoic gold deposits in the Japanese Islands. 1: Quartz vein-typegold deposits with little to no base metals. 2: Gold silver deposits with abundant base metals. 3: Distribution boundary of gold deposits formed during the Miocene. 4: Location of PlioPleistocene gold deposits at the actual island arc junctions. 5: Location of Plio-Pleistocenegold deposits in front of the actual island arc junctions. Numbers in the figure are K-Ar ages of epithermal Au-Ag veins (Kubota, 1994). Several sub-classifications of epithermal precious-metal deposits have been proposed: mineralogy, host-rock composition and elemental association (Lindgren, 1928), gold-silver ratios of metal weights (Ferguson, 1929; Nolan, 1933), mineral paragenesis (Nishiwaki et al., 1971), and production ratios of metals (Heald-Wetlaufer et al., 1983). Recently, epithermal gold deposits were divided into several types based on gangue minerals, and physicochemical environment of ore deposition (pH, H2S concentration of ore fluids). They are hot spring-type (Silberman, 1982; Berger, 1983a; Berger and Eimon,
Miocene-Pliocene Hydrothermal Ore Deposits
~'~A/kj~'~'+ " +~.~. L'ff~.~
.
~.
'.
J
.
~
~ .
v
~
ALTERATION
~XZ2~%8~\~,// ~ # \\%//..~.~,// ~/!
~. .
~
~,~//--,~.Stock. It Sill. . . . . Ore
AlteredB .
~,o~,p,,~
//-~ Ferruginous Chert (Fe-Mn) Barite Ore Black Ore ~ YellowOre ~
Epidote-rich Basalt 300m Acid/stuff'/,,~, ~,..--,/ ~ / ~ "
~ - ~. ..' .. ..~. ,',
11
C h l o r i t e - r i c h Basalt
O ~~" ~
+~'Y++'I:~&i ~ ,. .. .. .. .. .. .. . . ....... ~..:::iii ,o~,o
. . : . : . : ~ . ~ , ,
Clay/Mudstone Sericite/Chlorite Alteration
,~l
400m
Figure 1.7. Schematic distribution of Kuroko orebody and hydrothermally altered rocks (modified after T. Sato, 1974).
Shishimano Dacite L~
Age 0.66 l.lMa
300 , /<'/
Portal/', t
~ ~ &'~
N / \ Hishikari Lower t Andesites \
," ", - ,, ,- -
Old Hishikari'~"~"~.~• \ / / '~ ,. ,, -Yamada / ~ \
\t \ /\/'" "~ \ 1' N / \ ,
d W \/\ x,"/ x\,'\ ," \ ? ' 4 / - ~ v V ~ ' - - < ) , l ~ / \ , 0.95 ;" ,,, z \ ~ / x/\ / \ ,,"%.v,.~jc.,,v~7"-ff-~
100ML'~-,,.kL.." ~ ; - . ; ~ { , / , .
r/=/ =I~I
i~, '~
x'-"--x"-/Jf" z ""
~
I
1r
Vein
(Honko)
/
/
~-~
v
\ " \ z \ z \"~'4 1"79Ma
~
/
tl (Sanjin)
,,,
\,,
pre -Neogen~
0.78-1.05Ma Figure 1.8. Schematic northwest-trending section across the Main and the Sanjin deposits of Hishikari mine (Ibaraki and Suzuki, 1993).
12
Chapter 1
TABLE 1.2 Estimated total productions of Au, Ag and other metals and Ag/Au total production ratio (Ag/Au, by weight ratio) from the individual vein-type and disseminated-type deposits in Japan (Shikazono, 1986). Type I-A: gold-silver-rich deposits, Type I-B: base-metal-rich deposits, Type 2: disseminated-type deposits Mine
Deposit type
Au (M.T.)
Ag (M.T.)
Ag/Au
Hokuryu Sanru Numanoue Kohnomai Kitami Tokusei Taiho Teine Oe-Inakuraishi
l-A I-A I-A 1-B I-B 1-A I-A I-B 1-B
2.9 6.7 1.1 71.4 0.2 1.2 0.12 10.2 2.5
11.3 40 81.2 1219 22.2 14.7 10.2 158 109.5
3.9 6.0 71.4 17.0 11I 12.0 85 15.5 43.8
Todoroki Toyoha
1-A 1-B
5.7 2.3
209 914
37 404
Eniwa Koryu Chitose Oogane Shizukari Yagumo
1-A I-A I-A 1-A 1-A 1-B
0.71 0.76 22.8 1.6 7.4 0.4
5. I 22.2 105 35.0 5.1 70
7.1 29.4 4.7 19.4 7 175
Jokoku
1-B
-
152
-
Furokura
I-B
1.2
45
37.5
Osarizawa
l -B
6.2
251
40.5
Ani Takanosu Matsuaka
I-B I-B 1-B
1.0 0.6 1.53
32.0 3.9 62.6
30.8 6.6 40.9
Innai Hosokura
1-A l -B
1.0 2.9
400 527
400 184
Isobekoyama Handa Yatani
1-B 1-A 1-B
3.2 1 1.7
2.3 13 64
0.73 13 39
Other metals (M.T.)
Cu: 4720 Pb: 8850
Cu: 7291 Cu: I924 Pb: 17316 Zn:48100 Mn:307840 Pb: 226410 Zn: 558478
Cu: 70
Pb: 12000 Zn: 23200 Mn: 6444 Pb: 13810 Zn: 35906 Mn: 2762 Cu: 15500 Zn: 262500 Cu: 341000 Pb: 806000 Zn: 155000 Cu: 20770 Cu: 4770 Cu: 3150 Pb: 25650 Zn: 39600 Pb: 21534 Zn: 588596 Cu: 3850 Cu: 1270 Pb: 29210 Zn: 58420
Years of production
1928-1943 1925-1974 1923-1951 1917-1974 1934-1964 1930-1942 1912-1928 1932-i971 1890-1974
1903-1974 i914-1974 1929-1943 1903-1957 1936-1974 1932-1955 1918 1962 1931-1969
1941-1978
1904-I974
1885 1931-1969 1908-1950
1871 - 1953 1898-1977 1932 1919-1966 1870-1974
Miocene-Pliocene Hydrothermal Ore Deposits
13
TABLE 1.2 (continued) Mine
Deposit type
Au (M.T.)
Ag (M.T.)
Ag/Au
Other metals (M.T.)
Years of production
Mikawa
1-B
1.7
29.4
17.4
1942-1961
Sado Takatama Takahata Nebazawa Tochigi Ashio Ohito Toi Seigoshi Mochikoshi Yugashima Rendaiji
1-A 1-A 1-A 1-A 1-A 1-B 1-A 1-A 1-A I-A 1-A 1-A
57-77 28.8 3.1 1.0 0.115 3.1 1.03 18.4 13.5 4.9 2.2 5.6
1310 279.9 2.9 65 7.6 390 2.36 214 455 104 29.8 276
Cu: Pb: Zn: Cu:
9.7 0.95 65 66 125.8 2.3 11.6 34 21 13.5 50
Nawaji Shimonomoto Kishu Okinoura Takeno Kohmori Nakase Ohmidani Ikuno
1-A 1-B I-B 1-A 1-A 1-B I-B 1-A 1-B
1.5 1.3 2.2 4.5 4.6 0.123 2.6 0.3 2.1
25 80 179 4.4 91 28.0 12.8 79 403
16.7 62.5 80 1.0 19.5 228 4.9 267 194
Tada
1-B
0.0
0.2
8000
Sakoshi-Odomari OmorI Bajo Talo Fuke Okuchi Onoyama Yamaganc lsobe-Arakawa Kushikino Akeshi Kasuga Iwato
1-A 1-B 1-A 1-A 1-A 1-A 1-A 1-A 1-A 1-A 2 2 2
1.1 1.4 13.0 36.6 1.9 21 1.3 37.1 1.5 53.6 2.4 3.3 4.4
9.7 65.7
8.8 48.1
158.6 1.1 15.6 0.6 37.1 10 488 1.4 0.69 5.7
4.3 0.6 0.74 0.43 1.0 6.7 9.1 0.6 0.21 1.3
5281 1686 7857 5400
Cu: 671795
Cu: 1000 Mn: 15840 Pb: 3680 Cu: 89436
Cu: 20910 Sb: 1921 Cu: Pb: Zn: Sn: Cu: Pb: Zn:
76076 27664 152152 1521 12 2 7
Cu: 6331
1601-1970 I429-1974 1930-1976 1942-1974 1908-1950 1877-1966 1930-1952 1916-1965 1935-1976 1929-1962 1937-1972 1914-1959 1929 1971 1956-1962 1940-1974 1925-1942 1920-1949 1928-1968 1956-1966 1914-1974 1940-1973
1940-1973
1977-1982 I889-1918 -1945 1903-1973 1937 1947 1905-I974 1934-1963 i628-1955 -1970 1914-1974 1915-1974 1929-1966 1932-1980
Chapter 1
14
Cu
0
"Y" sub-type 0
0.- ~"~"~" .~s S @@
"C" sub-type
"B" sub-type
Pb
~@
•
8
•• •
\ Zn
Figure 1.9. Available data on the Cu, Pb and Zn ratio of total ore in a single unit deposit in the HanaokaKosaka district, marked with three sub-typesof Kurokodeposits (Horikoshi and Shikazono, 1978).
1983), quartz-alunite-type (Berger, 1983b; Berger and Eimon, 1983), high sulfidation and low sulfidation-type (Hedenquist, 1987), a low and high sulfur distribution and an alkali-type (Bonham, 1984, 1986), and Te- and Se-bearing types (Shikazono et al., 1990). Most of epithermal precious-metal vein-type deposits in Japan can be classed as adularia-sericite-type, and low sulfidation-type. Very few hot spring-type deposits (quartz-alunite-type, high sulfidation-type) are found in the Japanese Islands. A summary of various characteristic features of adularia-sericite type (low sulfidation-type) is given mainly in section 1.4. A few examples of hot spring-type deposits occur in the Japanese Islands. The characteristics of this type of deposits are described briefly in section 2.7. Shikazono et al. (1990) divided epithermal precious-metal vein-type deposits into Te-bearing and Se-bearing deposits. As will be considered later, Te-bearing deposits are regarded as intermediate-type of adularia-sericite-type and hot spring-type. The distinction between these two types of deposits is discussed in section 1.4. During the Miocene age, polymetallic vein-type (xenothermal-type, subvolcanictype) and gold-quartz vein-type (mesothermal-hypothermal-type) mineralizations occurred mainly in middle to western part of Japan. They are described in section 1.6.1. In section 1.6.2, Hg and Sb vein-type deposits are described. Each deposit type is distributed in a different metallogenic province (Fig. 1.3) (Tatsumi, 1970). Epithermal vein-type deposits occur in Miocene-Pliocene volcanic terrain.
Miocene-Pliocene Hydrothermal Ore Deposits
15
Polymetallic vein-type deposits occur in middle Miocene volcanic terrain in central and western Japan.
1.3. Kuroko deposits Hirabayashi (1907) defined "Kuroko" as an ore which is a fine compact mixture of sphalerite, galena, and barite. This definition can be applied to "black ore", but not to "yellow ore" or "siliceous ore" because these minerals are not abundant in these ores. Kinoshita (1944) defined "Kuroko deposit" as a deposit genetically related to the Tertiary volcanic rocks, consisting of a combination of Kuroko (black ore), Oko (yellow ore), Keiko (siliceous ore), and/or Sekkoko (gypsum ore) (Matsukuma and Horikoshi, 1970). The deposit is generally defined as a strata-bound polymetallic sulfide-sulfate deposit genetically related to Miocene bimodal (felsic-basaltic) volcanism (T. Sato, 1974).
1.3.1. Geological characteristics 1.3.1.1. Distribution Kuroko deposits occur in the Green tuff region which is characterized by thick altered volcanic and sedimentary piles of Miocene age. Distributions of Kuroko deposits and names of the representative mines are given in Fig. 1.10 and Table 1.1. Metals produced during the past are summarized in Table 1.1. Large Kuroko deposits occur in the Hokuroku district in Northeast Honshu (Fig. 1.11). Small numbers of Kuroko deposits are found in other districts such as Southwest Hokkaido, the northern part of Honshu (Shimokita Peninsula district), the middle of Northeast Honshu (Wagaomono and Aizu districts) and western Honshu (San-in district) (Fig. 1.10). It is clear in Fig. 1.10 that the distribution of Kuroko deposits is restricted in a narrow zone in the Green tuff region which was called a "Kuroko belt" by Inoue (t 969). This belt was formed by rapid subsidence under the extensional stress regime and is thought to have been a back-arc depression zone at middle Miocene age. The relationship between tectonic setting and formation of Kuroko deposits is discussed in section 1.5. 1.3.1.2. General geology, country rocks and tectonic setting A large number of studies on the general geology and stratigraphy in the Kuroko mine areas have been done. During t960-1970 many drillings were carried out by metal mining companies and the Metal Mining Agency of Japan. These data clarified geologic structure and stratigraphy of the mine areas. Many Kuroko deposits are distributed in the Hokuroku district, Northeast Honshu (Fig. 1.9). Therefore, general geology and stratigraphy of the Hokuroku district is briefly described below mainly following T. Sato (1974, 1977), Tanimura et al. (1983) and Ishikawa (1991) (Table 1.3). The lowest rock units are basement rocks, composed of phyllites, cherts, and minor sandstone probably of Paleozoic age. The oldest Tertiary formation, which is called Ohya
Chapter 1
16 oo
HOKKAIDO 200 km
HOKUROKU DISTRICT WAGAOMONO
04
DISTRICT
AIZU
HONSHU SANIN
~. ~' <~ S~~HIKOKU
~
~.~ /
JKYUSHU IG GI ( F
Q
~
~t4
DISTRIBUTION OF KUROK0-TYPE MASSIVE SULFIDEDEPOSITS IN JAPAN °.% Network, Stratiform sulfide, Gold pyrite, * °%
~,~
Pyrite, or typical Kuroko
Kuroko-typeGypsumor Barite Green TuffBeltof Japan
f;3~!~.~'~ Clusters of Kuroko Deposits
Figure 1.10. The distribution of the Green Tuff belt of Japan and the Kuroko-type massive sulfide deposits within it. Major mining districts are labeled and ore deposit clusters outlined (Cathles, 1983a).
(Menaichizawa, Sasahata), is composed mostly of brecciated andesite lavas and andesitic hyaloclastics. This formation is conformably overlain by the formation (Hotakizawa, Sunakobuchi) composed of thick sequence of basaltic lavas and tuff breccias with minor intercalations of mudstone and felsic tuff. The formation which is mostly composed of dacite lavas, tuff breccia and mudstone (Hanaoka, Yukisawa, Uwamuki formations) conformably overlies the Hotakizawa and Sasahata formations. The thickness of these formations is 300-400 m. Kuroko ore deposits occur at the upper part of this formation. White rhyolite lava domes characterized by intense sericite alteration have a close spatial relationship with Kuroko deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
N
|
~
x
x,z
,--'"
I /~
/
",,)
X~,.
/.-
Oinzan~, ~LakeTow..~ada
? /CNamanyamaI~ Furutobe ,~,. ~ ~-
/~ H~naoka! ].:/ / d~ IShakanm"/
"I
__
17
\-~'x '~Kosaka'.,~ I , _):.-~'~- ~'lt /~" " ~
-
',,/,....
r=k===W=~.-,:.\ \ 1 ~-
J
.~Somaki i" " ",~
; ( ( ./'=.... -4"
Osarizawa ~, ',~,,=
~"
•
/ t
'/~fd ]
.,.
"-,.,
O |
-'--"~ Fault ----'-1 Anticline [--'~---1 Syncline
~ ~L. ~
M. Mioc. Sed. Basin
Mioc.
Sed. Basin
10 km i
|
---] Kuroko Deposit ~
Network Vein Deposit
Principal
---1 Vein Deposit lntrusives Figure 1.11. Geologic structure and ore deposits in the Hokurokudistrict, Akita Prefecture,North Honshu (T. Sato, I974).
The formation composed of alternation of dacitic tuff and mudstone and basalt lava (Tsutsumizawa, Kagaya and Akamori formations) conformably overlies the Hanaoka, Yukisawa and Uwamuki formations. These formations are on average 150 m in thickness. The interbedded felsic tuff and mudstones (Shishigamori, Shigenai and Harukizawa formations) of middle Miocene conformably overlie the Tsutsumizawa, Kagoya and Akamori formations. The thickness of these formations varies widely, ranging from 40 m to 250 m. The younger formations (Ittori, and Tobe formations) of late Miocene to Pliocene overlie the Shishigamori, Shigenai and Harukizawa formations and are comprised mostly of mudstones, interbedded felsic tufts, and tuffaceous sandstones. The total thickness of these formations is ca. 500 m. The formations of Pleistocene unconformably overlie the
18
Chapter I
TABLE 2.3 Simplified stratigraphic column in the Hokuroku district and correlations with the Oga stratigraphy (Tanimura et al., 1983) HOKUROKU DISTRICT AGE
BLOW'S ! OGA PENINSULA M.Y ZONE
SHAKANA]MINES CO.,LTO. and AK{TA PRER GOV~
o
PLEISTOCENE -
- 5-
LATE MIOCENE
N.20 N.I9 N.)8
FunokowoE
N.I6 N.IS
Sqinokomi E (01akI F.)
*A-z~.
~-
MO-b rO-b MI-a TI
Ul-b T2182
N{ NJ N.[1
Shokona{ F. (Okuzu E)
N.S
~aolln S,dti.
_~;
TO-o
Itl6ri E
OnnogawoE (marine)
~
Tobe E
(marine)
Ealtlr,B TowedoVo~nlCl
T,.oo,o,~,.,J
BGztn~ MO-a'
k.14
MIDDLE MIOCENE
Central
J~IU¥1Um~
K]lauro I~ (marine)
N.17
IO
OOWA MININB CO., LTD.
Weslern
JI/T-W
N.2~ N.ZZ
PLIOCENE
M.M.A.j.
EARLY MIOCENE
{
~
"-~,TIUIlU~II¢IG"F'.'.-'~'. XogayeE~-,*.-.-.- Akofaorl F,o- %~
..... , HoaooknF..~. . . . . . .
Uwomukl
Su~okobuchl P
E ", : ;-~
~
T5/M3/B~
~20-
Yuk~snwa
T4/04
"
l:
'
"' "
. . . .
Ho~ok~zawa~
"
"
'
"
*.
'
'
"
. . . . .
"
~
"'"~
OLIGOCENE 25"
"60" /
MonzenG.
Ohya F. (MenoLchlzawa F.}
Toy/Ts
Oa~.ement
9aslme~!
P
~,,~;c:,lzewa F. : : : : : ' : : : : : : : : : ~~,:::::::::::::::::::::::::::::::::::::::::.::
L':::':::::::::'
Right column represents the average thicknesses and names of formations in three parts (western, central, and eastern) of the Hokuroku district, adopted by the Dowa Mining Company, Ltd.; second and, third columns from right are, respectively, the formation names adopted by the Metal Mining Agency of Japan and by the Shakanai Mines Company, Ltd., and the Akita Prefecture government; heavy horizontal and near-horizontal lines represent specific times, for example, the Menaichizawa and Sasahata formations are contemporaneous and correlate with parts of the Monzen Group and Daijima Formation at Oga.
a b o v e formations. T h e formations are basin-fill sediments, Towada v o l c a n i c s and terrace and a l l u v i u m deposits. T h e g e o l o g i c history in the H o k u r o k u district is divided into five phases as follows (Tanimura et al., 1983). ( l ) A n d e s i t e v o l c a n i s m and subsidence from near-shore to shallow-sea environments ( 3 0 ( ? ) - c a . 17 m.y. ago). (2) M a j o r subsidence, formation o f the H o k u r o k u basin and basaltic v o l c a n i s m ( 1 7 - 1 6 m.y. ago). (3) Felsic and basaltic v o l c a n i s m s and f o r m a t i o n o f K u r o k o ores ( 1 5 - 1 4 m.y. ago). (4) A c c u m u l a t i o n o f m u d s t o n e s and tufts in q u i e s c e n t sea ( 1 4 - 5 m.y. ago).
Miocene-Pliocene Hydrothermal Ore Deposits
19
(5) Differential uplifts, broad folding, andesitic volcanism and formation of smaller basins (5 m.y. ago-present). The most important geologic events at the time of Kuroko mineralizations are rapid subsidence just before the mineralization and bimodal volcanism (contemporaneous basic and felsic volcanism) (Konda, 1974). Several different hypotheses on the tectonic setting of the Kuroko mine area have been proposed. They include volcanic front of island arc (T. Sato, 1974; Horikoshi, 1975a), rifting of island arc (Cathles, 1983a), back-arc depression (Fujioka, 1983; Uyeda, 1983), and back-arc basin. In recent years, many hydrothermal solution venting and sulfide-sulfate precipitations have been discovered on the seafloor of back-arc basins and island arcs (e.g., Ishibashi and Urabe, 1995) (section 2.3). Therefore, it is widely accepted that the most Kuroko deposits have formed at back-arc basin, related to the rapid opening of the Japan Sea (Horikoshi, 1990). The summary of the bulk chemical compositions (major elements, minor elements, rare earth elements), 87Sr/86Sr (Farrell et al., 1978; Farrell and Holland, 1983), microscopic observation, and chemistry of spinel of unaltered basalt clarifies the tectonic setting of Kuroko deposits. Based on the geochemical data on the selected basalt samples which suffered very weak alteration, it can be pointed out that the basalt that erupted almost contemporaneously with the Kuroko mineralization was BABB (back-arc basin basalt) with geochemical features of which are intermediate between Island arc tholeiite and N-type MORB. This clearly supports the theory that Kuroko deposits formed at back-arc basin at middle Miocene age.
1.3.1.3. Age of mineralization The age of Kuroko mineralization can be estimated from (1) K - A r ages of igneous rocks associated with Kuroko deposits and (2) foraminiferal assemblages in mudstone directly overlying Kuroko deposits. Many K - A r ages of igneous rocks in Kuroko mine area in Hokuroku district yield 11-16 Ma. Ohmoto (1983) considered based on these data that the age of Kuroko formation was 11-16 Ma. However, almost all of igneous rocks associated with Kuroko deposits were hydrothermally altered after the mineralization (Shikazono et al., 1998). Therefore, it is likely that the K - A t ages of altered igneous rocks give younger ages than that of Kuroko mineralization. In contrast to K - A r ages, foraminiferal ages yield more accurate ones. Horikoshi (1987) investigated the foraminifera1 assemblage in mudstone overlying Kuroko deposits in the Shakanai area of Hokuroku district and concluded that the Kuroko mineralization age is 15-16 Ma. However, foraminiferal age data obtained by Horikoshi (1987) are scarce. Yoshida and Yamada (2001) compiled K - A r age of igneous rocks in Hokuroku district (which are considered to have formed simultaneously with Kuroko deposits) at 12.7 Ma. Considering uncertainty of foraminiferal and K - A r ages it seems reasonable that the Kuroko deposits in Hokuroku district formed in 14-12 Ma (more likely 13.612.7 Ma).
20
Chapter 1
However, several small Kuroko deposits (e.g., Yunosawa in Hokuroku district, Kuroko deposits in Hokkaido) occur in the formation younger than middle Miocene age (16-14 Ma), suggesting younger ages (12-13 Ma). 1.3.1.4. Metals enriched and metal ratios Many elements are concentrated in Kuroko deposits. Metals and minerals recovered from the ores are Cu, Pb, Zn, Fe, Au, Ag, S, BaSO4 and CaSO4. Size, average ore grade and tonnage of representative ore deposits are summarized in Table 1.1 (Tanimura et al., 1983). Total tonnage concentrated in the Hokuroku district is Cu, 1.2 x 108 ton; Zn, 2.5 x 106 ton; and Pb, 0.8 × 106 ton, respectively. Tonnage of BaSO4 and CaSO4 is given in Table 1.4. Figure 1.9 shows the proportion of Cu, Zn and Pb contents of Kuroko ore (Tatsumi and Ohshima, 1966; Horikoshi and Shikazono, 1978). Horikoshi and Shikazono (1978) divided Kuroko deposits in the Hanaoka-Kosaka area of Hokuroku district into three sub-types based on the ratio of Cu to Pb and Zn which increases in order of the B (black ore), C (composite ore), and Y (yellow ore) sub-types (Fig. 1.9). Characteristic features of these three sub-types were summarized by Horikoshi and Shikazono (1978) and are briefly decribed below. The Shakanai No. 1 deposit in the Shakanai mine is a good example of the B subtype (Kajiwara, 1970a). The ore of the B sub-type deposits consists of predominantly of galena and sphalerite with lesser amounts of chalcopyrite. Ore deposits of this sub-type are usually not directly associated with dacite lava dome. However, it is known that domeshaped dacite occurs below some of this sub-type deposit (Kajiwara, 1970a; Tanimura et al., 1974). Total ore quantity of a single unit deposit is generally small, about one million tons. C sub-type deposits are often called typical Kuroko deposits. Sato (1970) and Horikoshi (1976) published the schematic sections of Kuroko deposits referring to the general geology of this sub-type. The major Kuroko deposits belong to this sub-type. The largest Kuroko deposit is the Doyashiki deposit in the Hanaoka mine which belongs to C sub-type. The total ore quantity may be more than 10 million tons. The second largest deposit of Kuroko deposits is the Motoyama deposit of this sub-type. About 7 million tons of
TABLE 1.4 Estimated total amount of barite and sekko (gypsum + anhydrite) Shikazono, t983) BaSO4 (105 t) Kosaka
Fukazawa
Uchinotai Uwamuki Motoyama Tsunokakezawa Manjaku Kanayamazawa Matsumine
13. l 5.7 I 1.8 10.6 2.6 1.5 12
CaSO4 (105 t) Wanibuchi Fukazawa Yokota Motoyama Hamago
16 1-2 14 10
Miocene-Pliocene Hydrothermal Ore Deposits
21
ore were mined. The ore contains 2.3 million tons of black ore, 1.1 million tons of yellow ore and 3.7 million tons of siliceous ore. This means that ore deposit consists of roughly equal quantities of black, yellow and siliceous ores. A part of geology of the C sub-type is exhibited strikingly in the abandoned open-pit of the Motoyama deposit. The Uchinotainishi deposit of the C sub-type was described by Horikoshi (1969). The hydrothermal activity responsible for the mineralization of the C sub-type deposits was mostly preceded by the uplift of lava dome and the subsequent steam explosion (Horikoshi, 1969). Some of the Kuroko deposits consist predominantly of pyrite containing a small amount of chalcopyrite. The ore deposits consisting predominantly of pyrite, either with an economical value of chalcopyrite or not, are called the Y sub-type deposits, which occur above dacite lava dome or lava flow, while copper-poor deposits occur mostly in pyroclastic rocks and are associated with a large amount of gypsum. The Matsumine deposit in the Hanaoka mine is typical of the Y sub-type. The Matsuki and Takadate deposits in the Matsuki mine are also classed as this sub-type (Kuroda, 1978). Many pyrite-rich ore bodies
1oJG>. t omk2
°m/ m
~
~ U c h i n o t a i - n i U
-200m
-204~ •/
\
m 0
-100 m
-200 m
Figure 1.12. Distribution of two different sub-types of Kuroko deposits in the Kosaka district, Akita Prefecture. "Y" sub-type deposits have not yet been discovered in the area. The top pre-Tertiary basement is contoured showing some depressed structures (Horikoshi and Shikazono, 1978).
Chapter 1
22
associated with a large amount o f g y p s u m ore were mined in the Hanaoka mine. It seems likely that these ore bodies are composed o f several unit deposits o f this sub-type. Figure 1.12 shows the areal distribution o f the B and C sub-type deposits in the Kosaka district. The Y sub-type deposits have not yet been found in the district. It appears that two zones characterized by the distribution of each sub-type deposit are distributed north-southernly in the Kosaka district as well as in the Hanaoka district (Fig. 1.13). Pyroclastic rocks in the Kosaka formation, in which all deposits occur, become thicker to the east, and probably moved from the eruptive centres to the east (Horikoshi, 1969). These types of evidence may indicate that the sea at that time became deeper to the east. Figure 1.12 shows also the top of the pre-Tertiary basements. Ore deposits, either B or C sub-type, occur above the crater-like depressions o f basements. The Shinsawa deposit is the sole example o f B sub-type in the midst of the H a n a o k a - K o s a k a district, so-called Hokuroku basin (Fig. 1.13). The Tsunokakezawa deposit in the Fukazawa mine and ore deposit in the Ezuri mine are also the B sub-type.
I
.( \
/,Ioo..
g ; ',,.l_..~"
Tsutsumlzawo
\
J
N
Ok
X .\_,oo.
:\
',,x
/-"
m
Bsub,ype
"C" sub-type
\ -.,oo,,
/
"/~ ~
~
J
G ShokooaiNo,tf
0
1000 m
Figure 1.13. Distribution of three different sub-types of the Kuroko deposits in the Hanaoka district. The top of MI mudstone is also shown to visualize the structure of country rocks (Horikoshi and Shikazono, 1978).
Miocene-Pliocene HydrothermaI Ore Deposits
23
The Y, C and B sub-types roughly correspond to types 1, 2 and 3 as defined by Urabe (1974a), who classified Kuroko deposits based on hydrothermal alteration and ore mineral assemblages: type 1, kaotinite-pyrophyllite~diaspore-type; type 2, sericitechlorite-type; type 3, sericite-chlorite-carbonate-type. Hydrothermal alterations in the Kuroko mine area are described in section 1.3.2. Most large Kuroko deposits belong to type 2 (or C-subtype). Type 1 occurs mostly in Northeast Honshu and Hokkaido. Type 3 deposits are distributed in middle Honshu and Southwest Honshu (San-in district). Most of the previous studies have been carried out on the deposits in the Hokuroku district. A summary of the mineralogical and geochemical characteristic features of Kuroko deposits in this district is given below (sections 1.3.2 and 1.3.3).
1.3.2. Mineralogical characteristics
1.3.2.1. Metal zoning, and ore and gangue minerals
Typical Kuroko deposits (C sub-type according to Horikoshi and Shikazono (1978)) are usually composed of gypsum ore, siliceous ore, yellow ore, black ore, barite ore and ferruginous chert ore in stratigraphically ascending order (Fig. 1.7). The main constituent minerals in each ore are as follows: gypsum, anhydrite, Mg-chlorite (gypsum ore), quartz, pyrite, chalcopyrite (yellow ore), sphalerite, galena, pyrite, barite, chalcopyrite, tetrahedrite-tennantite, bornite, electrum (black ore), barite, quartz (barite ore), microcrystalline quartz, hematite (ferruginous chert ore). Figure 1.14 shows the distribution of minerals in each ore zone (Matsukuma and Horikoshi, 1970). The occurrence of ore minerals in Kuroko deposits was described in Shimazaki (1974), Matsukuma et al. (1974) and Urabe (1974a). Sphalerite has been studied by many workers and has been used to restrict the chemical environment of ore deposition. Iron contents of sphalerite from replacement siliceous ores and fissure-filling vein ores are much higher than those from layered ores (black ore and yellow ore) (Takahashi, 1963; Sato, 1969; Urabe, 1974b; Urabe and Sato, 1978). That is to say, iron contents of sphalerite generally decrease stratigraphically upwards in a single unit of ore deposit (Urabe and Sato, 1978). Urabe (1974b) and Urabe and Sato (1978) explained this trend by increasing of oxygen fugacity due to the mixing of hydrothermal solution with ambient cold oxygenated seawater. However, it is also likely that this trend was caused by the decreasing of temperature towards stratigraphically upwards. Generally, iron contents of sphalerite buffered by iron silicates such as chlorite decrease with decreasing of temperature in active geothermal systems (Fig. 1.15) (Hayba et al., 1985). Chlorite is common gangue and alteration minerals in Kuroko deposits. Therefore, this process seems plausible. Iron contents of sphalerite are different in layered ores in different ore deposit and different sub-types of Kuroko deposits. Iron contents of sphalerite from the B sub-type deposits (Uwamuki No. 4, Shakanai No. 1, Ezuri, and Fukazawa deposits) show wide range but generally less than 0.2 wt% (Ono and Sato, 1995). The average value, however, is probably lower than the C sub-type deposits (e.g., Uchinotai deposits). This may
Chapter 1
24
Mineral
Yellow ore Siliceousl ore Siliceous PowdeP/ Bedded PyCpSp ,to J YO YO E}O
Black ore .... GaFz Mononlne BO~nite BO raI~BO BO
Pyrite
Barite ore
Ferrugineus quartz
I
Chatcopyrite Bornits
u
Idaits
Fukuchilite Chalcocite
LII L
Covelline
,, ,
,L.
ass
Betlchtlnite Fehlerz
m
m m
Bournonite ,.
!,
., ,
Enargite Luzonite
immm mm
Argentite
l W
INm(m mmm me m m m m
Stromeyerits Jalpaite Polybasite Pyrareyrite Galena Sohalerite
IIIIIII i
m ~
III
Hi-disulphide
I
mm
!P
Mfflerite i iii
Bi,mlnerals iiiii
Electrum Silver
?
Sulphur Magnetite
4IlL
Hematite
i
Quartz I
III
I
Carbonates i ill
ii
III
Clays Barite
m
Figure 1.14. Schematic diagram showing mineralogical changes in various kinds of ores of Kuroko deposits (Matsukuma and Horikoshi, 1970).
suggest that the C sub-type deposits formed at higher temperatures and close to the volcanic centre and B sub-type deposits are distal type and formed at lower temperatures. Tetrahedrite-tennantite composition varies widely in Kuroko deposits (Yamaoka, 1969; Yamaoka and Nedachi, 1978a; Yui, 1971; Horii, 1971; Shimazaki, 1974; Kouda, 1977; Shikazono and Kouda, 1979; Ono and Sato, 1995; Ishizuka and Imai, 1998).
Miocene-Pliocene Hydrothermal Ore Deposits
25
0.5 ÷
0.4 0.3 0.2 03
0.1
14.
0 -0.1 0
,,-I
-0.2 -0.3 -0.4 -O.5 200
220
240
260
280
300
HomoganizationTemperature°C Fig. 1.15. Diagram showing the homogenization temperature of fluid inclusions vs. the iron content of the host sphalerite growth zone for sample locality NJP-X on the OH vein. The line shows the predicted iron content of the sphalerite if the sulfur fugacity of the system had been buffered by the triple point - - Fe-chlorite (daphnite), pyrite, hematite (Hayba et al., 1985). Generally, tetrahedrite-tennantite composition from Kuroko deposits is characterized by high Zn content, low Fe content, high Cu content, and low Ag content compared with those from vein-type deposits in Japan (Fig. 1.16). Rarely, it contains Hg up to 1 wt% (Ishizuka and Imai, 1998). Tetrahedrite-tennantite composition varies widely in a single orebody. For instance, Kouda (1977) analyzed tetrahedrite-tennantite from Fukazawa-Tsunokakezawa deposit and showed that the Fe and Zn contents are in a range of 0-5.5 wt% and 4 . 5 - 1 0 wt%, respectively. Wide compositional zoning and heterogeneity in a tetrahedrite-tennantite grain are c o m m o n (Yamaoka, 1969; Yui, 1971; Yamaoka and Nedachi, 1978a). Positive correlation between Zn contents of coexisting tetrahedrite-tennantite and sphalerite exists (Shikazono and Kouda, 1979). This relation can be explained in terms of the following reaction.
[(Cu, Ag)10Zn2(As, Sb)4Sl3]tet -]- (FeS)sp = [(Cu, Ag)10Fe2(As, Sb)4Sl3]tet +
(ZnS)sp (1-1)
where sp = sphalerite, and tet = tetrahedrite-tennantite. Electrum occurs mainly in the black ore zone dominantly composed of sphalerite, bornite, galena and barite. It is common in B sub-type such as the Ezuri and Fukazawa deposits. It occurs in brown compact black ore consisting of barite, galena,
Chapter 1
26 AG WT%
FE WT%
60
8
7
50
6 40
t'~: A
~
';',
A
5 4
30
3
20
2 10
1 f
10
20
30
40
1
50 CUWT%
2
3
4
5
6
7
8
9 10 11 12 ZN WT%
SB W'P/o 35
C
30 f..J t'. 25
D ",
A
,~
20
~ C ~'D
15
is
2o
:Is
AS WT%
Figure 1.I6. Chemical composition of tetrahedrite-tennantite (Shikazono and Kouda, 1979). A: Au-Ag vein-type deposits, B: Kuroko deposits, C: Taishu Shigekuma Pb-Zn vein-type deposits, D: Skarn deposits (Kamioka). sphalerite, tetrahedrite-tennantite, chalcopyrite, bornite, and Ge-bearing minerals (argyrodite), stromeyerite, pearceite, and mckinstryite (Ono and Sato, 1995; Ishizuka and Imai, 1998). Electrum is often associated with pearceite and/or tetrahedrite-tennantite (Ishizuka and Imai, 1998). Exceptionally, electrum occurs abundantly in the siliceous ore of the Nurukawa deposit (Yamada et al., 1987). The mode of occurrence of electrum was described by Matsukuma (1985). Chemical compositions of electrum from Kuroko deposits were summarized by Shikazono (1981) and Shikazono and Shimizu (1988a). The Ag content of electrum from Kuroko deposits varies widely from 4.7 to 89.4 atomic% (Fig. 1.17). Electrum with low Ag
Miocene-Pliocene Hydrothermal Ore Deposits
27
40
u
0"
30
LL
20
10
0
20
40
60
80
100
Nag Figure 1.17. Frequency histogram for the Ag content of electrum from Kuroko deposits in Japan (Shikazono and Shimizu, 1988b).
content occurs in siliceous and yellow ores. For instance, electrum in siliceous ore from the Nurukawa deposit contains about 20 Ag atomic%. Electrum with high Ag content occurs in black ore. Electrum in the brown ore occurring in upper part of black ore contains Hg up to 11 atomic% and Ag contents are positively correlated to Hg contents (Ishizuka and Imai, 1998). Native silver is found in bornite-rich black ore (Matsukuma and Yui, 1979; Matsukuma, 1985) and it is thought to be secondary mineral. Compositional zoning in electrum grain is common (Shimazaki, 1974; Imai et al., 1981). The Ag content of rim of electrum grain is higher than that of core. Although Ag content varies widely, it is generally lower than that of epithermal vein-type deposits. Although analytical data on bornite are few, some data show high Ag contents (max. 1.45 wt%) (Matsukuma, 1985). Pyrite is the most abundant ore mineral. It occurs as euhedral, framboidal, and colloform forms. Abundance of framboidal pyrite increases stratigraphically upwards. Colloform pyrite contains appreciable amounts of As and Cu (Nakata and Shikazono, unpublished), whereas these contents of euhedral and framboidal pyrite are less than the detection limit of an electron microprobe analyzer. Ishizuka and Imai (1998) found that the As content increases toward outer rim and reaches up to 5 wt% in the rim of colloform pyrite from the Fukazawa deposit.
Chapter 1
28 1.0.
~2:~?~~2 o' 002 $2
o
oI 03 of 04
0-8-
05
0"6" o6
oO6
=,
06 I.L
06
0606
O.Z,-
06
0-2
9oo-9 08 " ~ o 8 8 o8
o9 0"8
1~0
1'.2
1J4
o707 1:6
I ~8
2:0
A~ in 4 ( A I , S i )
Figure 1.18. Variation of Fe2+/(Fe2+ + Mg) and tetrahedral AI of chlorite from hydrothermal ore deposits: Japanese Neogene Cu-Pb-Zn vein-type (open circle) and Kuroko deposits (solid circle). Localities: 1 Ashio, 2 Yatani, 3 Toyoha, 4 Kishu, 5 Sayama, 6 Mikawa, 7 Furutobe, 8 Hanaoka, 9 Wanibuchi, 10 western Bergslagen (Shikazono and Kawahata, 1987). Dominant gangue minerals in Kuroko deposits are quartz, barite, anhydrite, gypsum, chlorite, sericite, and sericite/smectite. Morphology o f quartz changes from euhedral in the centre to the irregular in the margin of the deposits (Urabe, 1978). No amorphous silica and cristobalite have been found. Quartz is abundant in siliceous ore, barite ore and tetsusekiei ore. Minor amounts o f M g - m i n e r a l s (talc, M g - c h l o r i t e ) occur in sekko ore. Chlorite occurs in sekko ore and it contains high amounts of Mg (Fig. I. 18). Kuroko deposits are characterized by large amounts of sulfate minerals (barite, anhydrite, and gypsum). Estimated total amount of barite and sekko (gypsum ÷ anhydrite) from individual deposit is shown in Table 1.4. Sr contents o f gypsum, anhydrite and barite
Miocene-Pliocene Hydrothermal Ore Deposits
29
Matsumine
6O
c- 10-
<
0 ..Q
5
E Z .
.
.
.
.
500
1000
.
.
.
1500 Sr
2000
con tent
(ppm)
Shakanai
t~o
t'- 10
<
0 r.~
5
E Z
2o o Sr c o n t e n t (ppm)
Fukazawa
cO (l) cO >,
c- 10
<
0
$ ..Q
5
Z ....
N n .F1. ,
500
•
-
,
I000
•
.r]
-
,
.
.
.
.
t
~500 2000 Sr c o n t e n t (ppm)
Figure 1.19. Strontium contents of anhydrites from the Matsumine, Shakanai, and Fukazawa deposits (Shikazono et al., 1983).
are 150-2,000 ppm (gypsum, anhydrite), and 0.3-3 wt% (barite), respectively (Shikazono et al., 1983) (Fig. 1.19). Mixed layer clay mineral (sericite/smectite) is found in Kuroko ore bodies and altered dacitic rocks underlying the ore. This mineral is thought to have formed by the
30
Chapter 1
interaction of sericite with hydrothermal solution whose pH decreased by the deposition of sulfides by the reaction of MC12(aq) + H2S -+ MS + 2H + + 2C1- (where M is metal such as Zn and Pb) (Tamura, 1982). Tamura (1982) thought of the following mechanism of formation of mixed layer clay mineral (sericite/smectite). Mg 2+ and Ba 2+ in hydrothermal solution may play an important role to form the interstratification of the minerals. From the experiment in which Li + has caused the formation of the interstratified structure, Mg 2+ with most similar ionic radius (0.65 A) to Li + (0.60 ,~) may accelerate the interstratification. Tomita and Sudo (1981) and Shimoda et al. (1974) have synthesized the mixed layer minerals from sericite reacted with LiNO3 solution. Li + gets into the unoccupied site in the octahedral layer of dioctahedral muscovite (sericite). To retain electroneutrality in muscovite structure, interlayer K + is easily released. Potassium is held strongly adjacent to interlayer regions where the potassium is replaced, i.e., the two regions will alternate. This is interstratification. Reichenbach and Rich (1968) have revealed that interlayer K could be removed almost completely from muscovite by using 0.1 N BaC12 solution at 120°C. This K - B a exchange may induce to form the interstratification. To verify the above interpretation, the analysis of Ba and Li in sericite/smectite is required. Preliminary analysis of sericite/smectite shows high Ba content (Ogawa and Shikazono, unpublished).
1.3.2.2. Hydrothermal alteration Several studies identified the hydrothermal alteration halo in the dacitic rocks surrounding Kuroko deposits (Fig. 1.20) (Shirozu, 1974; Utada et al., 1974, 1981; Utada, 1980; Ishikawa et al., 1976; Izawa et al., 1978; Date et al., 1983; Urabe et al., 1983; Ishikawa, 1988; Marumo, 1989; Inoue and Utada, 1991; Shikazono et al., 1995). For example, Date et al. (1983) recognized the following alteration zones in the Fukazawa Kuroko mine area of Hokuroku district from the centre (near the orebody) to the margin: (1) sericite-chlorite zone (zone III in Figs. 1.20-1.22) characterized by quartz + sericite -t- Mg-rich chlorite; (2) montmorillonite zone (zone II in Fig. 1.20) characterized by Mg.Ca-type montmorillonite + quartz -t- kaolinite 4- calcite 4- sericite -tFe-rich chlorite; and (3) zeolite zone (zone I in Fig. 1.20) characterized by clinoptilolite + mordenite + Mg.Na-type montmorillonite 4- cristobalite 4- calcite or analcime + Mg.Na-type montmorillonite + quartz + calcite -t- sericite -t- Fe-rich chlorite (Fig. 1.20). Kaolin minerals (kaolinite, dickite, nacrite), pyrophyllite and mica-rich mica/smectite mixed layer mineral occur as envelopes around barite-sulfide ore bodies in the footwall alteration zones of the Minamishiraoi and Inarizawa deposits, northern part of Japan (south Hokkaido) (Marumo, 1989). Marumo (1989) considered from the phase relation in A1203-SiO2-H20 system that the hydrothermal alteration minerals in these deposits formed at relatively lower temperature and farther from the heat source than larger sulfide-sulfate deposits in the Hokuroku district. Date et al. (1983) found the existence of an Na20-depleted dacite mass with a lateral dimension of 1.5 x 3.0 km immediately below the ore horizon (Figs. 1.23 and 1.24) and the mass is useful indicator of exploration of Kuroko ore deposits. This Na20 depletion is considered to be due to the destruction of plagioclase attacked by potassium-
31
Miocene-Pliocene Hydrothermal Ore Deposits
TKIS;
/
I
N
TKI?9
f
D-3
I
o
I
TKt$3
®
/
/ I
/
/
\
/
e
TKII9
! I iK,ao
TKI78
•
t
\
zoNE,
0
e
\
J
/
DE 4
g
/I
;rK139 ~
I
,5~\ g~,oa
TK208 e
/
/ ZONE
II'l
tK3s III
/
1.0kin
TKI81 ®
/
/I 1
I
ZONE I
I I I
HO7
TKIgO 0
0.5
i
KA1
OES
I
TK177
TKt98
e
TK 20g e
\
®
\
~'~
\ \\\
o~~
TK200
e
'
I \ T~a,
\
LEGEND ORE
DEPOSITS
(3~i
."'''4
...........' ORE ZONE
~
, , ~ '~" - ' ~ " ~ .
T,.~.. TKt/~
~K,o3)
~"'
ALTERED DACITE I N FOOTWALL
....
•
ZONE I
4)
ZONE
II
O
ZONE It"
•
ZONE IU
•
ZONE
IV
TK207 e
TKI&~ e
TK 1~,6 e
ZONE BOUNDARY
Figure 1.20. Zoning map of alteration minerals in unit D3 around the Fukazawa deposits (Date et aI., 1983).
rich hydrothermal solution and the formation of K-sericite in discharge zone. It is thought that potassium is added from ascending hydrothermal solution and sodium in dacite removed to hydrothermal slution. Hashiguchi et al. (1983) also showed that the variation of Na20 content in the footwall rock is particularly useful for detailed exploration based on the large number of chemical analyses of footwall rocks and statistical treatment of the geochemical data. Singer and Kouda (1988, 1992) confirmed based on the statistical analysis of the distribution of minerals and bulk compositions of altered rocks in the Hokuroku district
Chapter 1
32 N
om
S
=
o
°T 1
-:7-
° T2
LEGEND
T1, MI.T2, M2 :ROCK UNITS IN HANGINGWALL
0.5
i
:ZONE I
~:ZONE
U ~nd II
~:ZONE
Ill
:INTRUSIVE OOLERITE
~:ZONE
IV
: BASALT LAVA
................ :ZONES BOUNOARY
:KUROKO DEPOSITS
~FF~:Na POORPART IN FOOTWALL
T3, D3, T4,04 :ROCK UNITS IN FOOTWALL :INTRUSIVE DACITE
{ ~
0
•
1,0 km
Figure 1.21. Zoning of alteration minerals in the Fukazawa area. The location of the profile line is shown in Fig. 1.20 (Date et al., 1983).
HANGINGWALL I u~ ~
ZEOLITE
11
It'
Ill
IV
ANALCI ME MOROENITE
MONTMORILLON I T E
Ca-MONT.
No-MONT.
0
==~ S~RICITE
~u
-~ i
0
CHLORITE
PLAG IOCLASE HYDROTHEIIMAL ALTERATION DI A G E N E ~ L,~
FOOTWALL I[ u') c= ~ t~ Q:
ZEOLITE
MORDEN I T E
NtONTMORILLON1TE
Co-MONT.
~
SERICITE
~
CHLORITE
III
Mg-rlch
PLAGIOCLASE HYDI !OTHERMAL
ALTERATION DIAGENES I S )EPLETIVE
ADDITIONAL
ELEMENTS
ELEMENTS
CoO
Na20 SiO2
CoO
No20 CoO, FeO
IMgO !FezO3
Figure 1.22. Summary of alteration minerals and zoning around the Fukazawa deposits (Date et al., 1983). Legend: unbroken line = present or formed in considerable quantity; dashed line = present in small amounts, or uncertain formation; definition of zones by Date et al. (1983).
Miocene-Pliocene Hydrothermal Ore Deposits
i
~FIAe..
"
/J~
. %.
I
;<-'~
~,
•
.
.'#
.~. ..~° •
•
"
. ~
o
~ o
°
" ~
~
.
"~j
o
.
.
o
o,/7/
"o
%
"
i
° o
o
*'//,<,t-a~
°
•
•
/ / Z Z J / ? a.
.
•
33
o
1".
•
~"/
o
°
•
"~..
~
o
.~.
°
- ~ "~;~//~.~7~,///~ ' ../
..,
/5"
• / .
__
o
~;-~,• "~.
".
°
~°
~-
"~.'~No%
o 0~.
o
•
~
•
°
~,o
i
o
o o
~'\ - ~ - ~ d . 0. . . .
~,,,.~o
~
• --
•
•
"~-2- FOoIw.II Expo .... (4 l
V
KUROKOOeposit
X /
,~lm01e,
/ /
t--7 I
Oril, hole
[_...v "7, token
o 13 or fewer sornples token
.~.'~ ~ % .
~-~r-.<~z.
~
• \
~, f
t
_"'~1 ~ > ' \ -
e~
~
,~:~.;'-. ..
o o
"
I No 5
~-,,g.,~./~%Z; e
x
/
/
°
• •
"'v06.--I~°.. I__/' o \ ~''Z-'~III/Xi .~./ '~'~
o
"
•'~Z'Z.'~';¢L/i° • \
o
" °
°
". 0100
soo
,ooo~
F i g u r e 1.23. Distribution m a p o f low N a 2 0 a n o m a l i e s in the K o s a k a a r e a ( D a t e et al., 1983).
that Na20 depletion, as well as the presence of sericite, anhydrite and gypsum near the orebody, is a useful indicator for the exploration of Kuroko deposits. Figures 1.25 and 1.26 show the map of probabilities of centres of Kuroko deposits based on the data by Singer and Kouda (1988). Green et al. (1983), Urabe et al. (1983) and Matsuhisa and Utada (1993) showed that the analysis of whole-rock 6180 values is a useful method of exploration for volcanogenic massive sulfide deposits. They found that concentric zoning patterns of the whole-rock 8180 values and of alteration minerals are particularly well developed in the footwall volcanic rocks in Kuroko mine area. For instance, 8180 values of igneous and sedimentary rocks are +16.9-4-2.7%o in the zeolite zone, +11.1+2.5%o in the montmorillonite zone, and +6.7 4- 1.3%o in the sericite-chlorite zone in the Fukazawa area (Green et al., 1983) (Figs. 1.27 and 1.28). Figure 1.29 shows the 6180 variations as functions of temperature and water/rock ratio (Green et al., 1983). This calculated result indicates that the spatial variation of 8180 values of the footwall rocks and alteration
Chapter I
34
. .
.
TWNOKAKEZAWA No2
.
$? /
LEGEND Low Na, 0 Zonr (Naz0<0. 3 2 1 )
SARUMA Vrln
KUROKO 0.porit Drill hole 15 or more w r n p l * ~ taken 0
./
0
.
/
4 /
!4 OT f.**,
sornpt..
y/q./. >.
token
M 0
100
500
1000 J
Figure 1.24. Distribution map of low Na20 anomalies in the Fukazawa area. The distribution of green dacite is also shown (Date et a]., 1983).
mineralogy can be interpreted as a result of interaction between the rocks and seawater at different temperatures (25-200°C for the montmorillonite zone and 200400°C for the sericite-chlorite zone) under water-dominated (e.g., waterlrock ratio > 1) conditions and most of the hydrogen isotope data of the whole-rock samples from the Hokuroku district (6D = -34%0 to -80%0) can be explained by this interpretation (Green et al., 1983). However, the variation of of altered rock depends not only on waterlrock ratio and temperature, but also on the degree of mixing of hydrothermal solution and cold seawater. This mechanism is considered in section 1.4.2. Ohmoto et al. (1983) found that 6 1 8 0 values alteration zones and major and base metal contents of footwall dacite along the section in the Fukazawa area correlate with each other (Fig. 1.30). Shikazono et al. (1998) found that carbonates are common alteration minerals in the Uwamuki mine area of Hokuroku district and carbonate alteration superimposed on chlorite alteration. They showed that the mode of occurrences and the Mg/(Mg
+
Miocene-Pliocene Hydrothermal Ore Deposits
I
N 5 km
Probability ~ ~7
P
[~
0.001<=P<=0.1
~
35
FURUTOBE
/
0.1
P>0.5
1
Kuroko Deposit Area considered ~
~
i
UCHINOTAI SHAKANAI
ODATE
TOWADA MINAMI
Figure 1.25. Map of probabilities of centres of Kuroko deposits based on sodium depletion, sericite, and gypsum plus anhydrite(Singerand Kouda, 1988).
Fe) ratios of magnesite and dolomite occurring in hanging wallrocks are useful in the exploration for concealed volcanogenic massive sulfide-sulfate deposits. 313C and 3180 of carbonates from the Kuroko mine area are plotted in Fig. 1.31. 313C and 3180 data lie between igneous (3~3C = -7%o, 3180 = +8%0) and marine carbonate value (313C = 0%0 and 3180 = +20%o). This indicates that magnesite and dolomite formed due to the interaction of hydrothermal solution with the biogenic marine carbonates. Dolomite and magnesite data are plotted close to marine carbonate values, suggesting that they formed in the central zone close to the ore bodies due to the interaction of hydrothermal solution with the biogenic marine carbonates (Fig. 1.31). 313C and 3180 of manganoan calcite in altered basalt directly overlying Kuroko orebody are close to igneous carbonate values, suggesting they formed from ascending hydrothermal solution at discharge zone.
36
Chapter 1
Figure 1.26. Map of probabilities of centres of Kuroko deposits based on sodium depletion in the Hokuroku district (Singer and Kouda, 1988).
Superimposed alterations are common in the Kuroko mine area (Inoue and Utada, 1991). For example, K-feldspar, kaolinite, alunite, pyrophyllite and diaspore alterations cut chlorite alteration, indicating that they formed later than chlorite alteration (Inoue and Utada, 1991). Inoue and Utada (1991 ) thought, based on detailed descriptions of the hydrothermal alterations in the Kamikita mine area, North Honshu, that hydrothermal alterations in this district started from 13 Ma and ended at 3 - 4 Ma. Pyrophyllite and diaspore alterations were reported from several Kuroko deposits, although they are not common (Urabe, 1974a). This type of hydrothermal alteration is thought to have occurred at a later stage than the hydrothermal alterations associated with Kuroko mineralization (sericite, chlorite, and zeolites) (Utada, personal communication, 1995). As well as felsic volcanic rocks, basalt occurs in the Kuroko mine area. It is also intensely and hydrothermally altered. Shikazono et al. (1995) studied the hydrothermal
Miocene-Pliocene Hydrothermal Ore Deposits
i/,
- N-
~11A.8
1"
,' e~
~'~
37
;
/ - 18.8
/
!/
I d 15-4
|
• | lS.7
~11::0
|
" ~ ~' "4
I~:~
I
I
'
~ ZONE I ] IL 13.3• -
~ 10,1 ~"11 6
.O ~ ~= ~ , , f 130
"~
SHINSAWA I°6s 78
•I
! 3.0
014.4
/
i _r~ 133" 11.8
f/ •• ~
[I
19.3 v.~17n .... 13.9
I
18.7
..2~2~. ..... I e14.3 11 ~FUKAZAWA t
~,
es: 3
ZOGAKURA ~l
1~.5 •
",~
!1o.8
,\
KANAYAMAZAWA ~
~ ~=='%1 lS.8 , ~ ~1o.7 TAKARAKURA t/[ III • t~z.8 ~,~ 58
\
...-'~',,~9 -
~4.6 5.6
LEGEND
8,5 010.5
I
~
~~8.8
SARUMA " /
''~
--A'
09.5
ZONE h Zeolite Zone Ih M o n t m o r i l l o n i t e Z o n e IIh S e r i c i t e - C h l o r i t e Z o n e
0 I
z
1 km I
: Kuroko Ore Body : Vein Deposits e6.0 : 8180 .R.(%o) Value
17.8 12.0
10.3 11.9
Figure 1.27. Areal distribution of the whole-rock8180 values of footwall volcanicrocks in the Fukazawa area. The boundaries for the alteration zones are modifiedfrom Date et al. (1983) (Green et al., 1983). alteration of basalt overlying the Kuroko orebody in the Furutobe mine area of Hokuroku district. They showed that hydrothermally altered basalt can be divided into chlorite-rich rock and epidote-rich rock. Chlorite-rich rock occurs widely, whereas epidote-rich rock occupies a smaller area, close to the orebody. It was found that the CaO, Na20 and SiO2 contents of the bulk rock correlate negatively to MgO content, while FeO and NFe contents correlate positively to MgO content (Fig. 1.32) and these changes can be explained by seawater-basalt interaction at elevated temperature. The MgO/FeO ratios of
Chapter 1
38
.
Mont.
Facies
[] Mudstone
Ser.+ Chl.
Basalt, Andesite
Facies
0
[] Tuff J • Dacite
5
10
15
20
25
8180 (%o) Figure 1.28. Whole-rock 3J80 values of Miocene volcanic and sedimentary rocks from the Hokuroku district, grouped by alteration zones. Each square represents one sample. Mont. = montmorillonite, Ser. = sericite, Chl. = chlorite, av. = average (Green et al., 1983).
chlorite and actinolite and the Fe203 content of epidote from the basalt are greater than those of midoceanic ridge basalt (Figs. 1.33 and 1.34) probably owing to the differences in the FezO3/FeO and M g O / F e O ratios of the parent rocks. The lower CaO content and the higher N a 2 0 content of the bulk rock compared with altered midoceanic ridge basalt are interpreted in terms of the difference in original bulk rock composition. 3D of epidote in hydrothermally altered basalt is in a range of -36.5%~ to 43.0%o (Shikazono et al., 1995). Using -37.5%0 as an average 3D value of epidote, the fractionation factor for the H isotope exchange reaction between epidote and H 2 0 (Graham et al., 1980), and 280°C as the temperature of formation of epidote estimated from fluid inclusion study was calculated to be +2.3%~. This value is close to that of hydrothermal fluid issuing from the East Pacific Rise 21°N (+2.0%~, Bowers and Taylor, 1985). This strongly suggests that epidote formed at a discharge zone in submarine hydrothermal system (Shikazono, 1984). It is also notable that the 3D values of epidote from the basalt from Kuroko mine area are close to those of hydrothermally altered midoceanic ridge basalt in the Costa Rica Rift and Galapagos Ridge (-31%~ to -45%~) (Kawahata et al., 1987). 1.3.3. Geochemical characteristics
A large number of geochemical studies on Kuroko deposits (fluid inclusions, gas fugacities, chemical and isotopic compositions of ore fluids etc.) have been carried out. These are summarized below.
Miocene-Pliocene Hydrothermal Ore Deposits 25
[6m0i.÷6-'
20
39
1
'/
'
(at
J
,
% 2~
1
I
4
'l
t
I
~
l
,
to
,
.
.
.
.
,
loo
.
m i 8 0~= 0%°
(b) Zeol. Zone
T=C---2 W e48,,~f 45 o UW.R.
200~, "t !::
5
....
~
.04
.1
2 5
I
'
r '
I
400
t0 '
'
'
,
I
.
.
.
.
(c)
¢:t8n f 15 u ~W.R.
(%.)
"r c-,-
£1
i
I
ii i
Water / Rock ( Atomic Oxygen) Figure 1.29. Calculated changes in the ~180 values of volcanic rocks (~I80 = +7.0%o) as a result of equiIibrium oxygen isotope exchange with waters of different initial compositions. The dotted areas represent the 3180 ranges of rocks in the zeolite and the sericite-chlorite zones (Green et al., 1983).
1.3.3.1. Fluid inclusions Fluid inclusions from Kuroko deposits were studied first by Tokunaga and Honma (1974) who showed that Kuroko deposits formed in a range of 200-260°C for the siliceous
Chapter 1
40
(N)I~3
]] Hole NO
114
180
lance (Kin} 4 .~ration Zone - . . ZeoI
161
I~8 %54~ 5
? ,
~
5~ I~68 1~1 I03
o Ore Zone 0 Mont
~,
{s}
z
Mont
20f
(t}]
(,.ol :(
1
(S)
(NI
Cu
i
I"
:
!
"
(ppm)
'PP" ':I'
'
~'"
'
KeO (wt.%) 4
Zn:f
mt
(ppm) Ioo MgO
(wt%)
CaO
:61
.
:t (wt%)
z
(wt.%)
OI
SF :t
{ppm)
I
0
Figure 1.30. Comparisons of the ~5180values of footwall dacite, alteration zones, and other geochemical halos along the section in the Fukazawa area (Ohmoto et al., 1983).
ore, 100-240°C for barite and 140-170°C for black and yellow ores. Lu (1969) reported homogenization temperatures of the fluid inclusions in quartz from the Uchinotai-Higashi deposit of the Kosaka mine to be 200-250°C. Watanabe (1970) studied fluid inclusions from Ainai deposits and found that the homogenization temperature for the siliceous ores (190-300°C) is higher than that for the black ores (120-290°C). Marutani and Takenouchi (1978)clarified the variations in homogenization temperature and salinity of inclusion fluids in quartz from stockwork siliceous orebodies at the Kosaka mine (Fig. 1.35; Urabe, 1978). They showed that the temperature decreases stratigraphically upwards from stockwork ore zone (280-320°C) to bedded ore zone (260-310°C). Pisutha-Arnond and Ohmoto (1983) carried out fluid inclusion studies of the stockwork siliceous ores from five Kuroko deposits (Kosaka, Fukazawa, Furutobe, Shakanai, and Matsumine) and revealed that black ore minerals (sphalerite, galena, barite) and yellow ore minerals (chalcopyrite, quartz) formed at 200-330°C and 330 4-50°C, respectively, and salinities of the ore fluids remained fairly constant at about 3.5-6 equivalent wt% NaC1. They analyzed fluids extracted from sulfides and quartz; Na ---0.604-0.16 (mol/kg H20), K = 0 . 0 8 ± 0 . 0 5 , Ca = 0.064-0.05, Mg ----0 . 0 1 3 ± 0 . 0 0 8 , C1 ---- 0.82-t-0.32, C (as CO2) -- 0 . 2 0 ± 0 . 1 5 and less than 6 ppm each for Cu, Pb, Zn and Fe.
Miocene-Pliocene Hydrothermal Ore Deposits
41
0 -1
-2
0
o ^~"~
~-3 0 ¢,0
//~///
.g/°--orJ..o" ,, 0///
+o0,% .~" . o
-4
/I
q. 0 /
0
-5 //
-6
~/
Q
0
o
0
-7 I
5
f
t
i
s
I
i
10
,
~
t
I
I
15 a18 0 ( ~ )
I
I
~
I
20
~
i
,
i
I
,
25
Figure 1.31. Bivariable plot of oxygen versus carbon isotopic compositions of carbonates. Solid circle: magnesite; open circle: dolomite; open square: calcite; A: oxygen and carbon isotopic compositions of igneous carbonates; B: oxygen and carbon isotopic compositions of marine carbonates (Shikazono et al., 1995).
Using homogenization temperature and freezing temperature data (Fig. 1.36) and pressure-temperature diagram of N a C 1 - H 2 0 - C O 2 system, minimum seawater depth at the time of Kuroko mineralization is estimated to be 1,000-2,000 m, if boiling of ore fluid did not occur (Fig. 1.37, Pisutha-Arnond and Ohmoto, 1983). This estimated depth is close to that of seawater associated with present-day hydrothermal mineralizations at back-arc basins such as Okinawa Trough (section 3.3). However, Lu (1983) found that the salinity and homogenization temperature of fluid inclusions from the Uchinotai-East ore deposit varies widely and thought that this variation is due to boiling of Kuroko ore fluids. If his argument was correct, the depth could be estimated to be 1,000-1,500 m. Two hypotheses of seaftoor depth at the time of mineralization have been proposed based on foraminiferal data, ca. 3500 m (Guber and Ohmoto, 1978; Guber and Merrill, 1983) and 1500 m (Kitazato, 1979). Considering seafloor depth of present-day ore formation at back-arc basins and fluid inclusion data mentioned above, shallow seaftoor depth hypothesis (Kitazato, 1979) seems more likely. If the pressure-temperature condition of Kuroko ore fluids was close to the boiling curve, the depth could be estimated to be 1,000-1,500 m, which is similar to that for present-day back-arc mineralization such as Okinawa Trough.
1.3.3.2. Gas fugacities Sulfur fugacity (fs2). If electrum is in equilibrium with argentite, the equilibrium constant for the sulfidation reaction (K1-2), 4 Ag (electrum) + 82 = 2 Ag2S (argentite)
(1-2)
where Ag (electrum) is the silver component in electrum, can be expressed as, KI-2 =
a2 I(a 4 ~ Ag2S/I, AgJS2)
(1-3)
Chapter 1
42 30"
O32
o26
10
I0 %o1~ o4
°z7 9
;i 0 • 6
023
~+
09
01
8
~e
01$
028
191 0 21 027
7
t$O 031
+
:ff
~,11
020 03005 08 01+,
O 6 0 (,..)
lO01~&
5
e2~
II 25
0++
7 ~$1
o23
2~ol 7
08
e
3
02 01 •
:+2om6
32
+o %
~3
)26
50 ols°3 12°
0%
2
o28 o~'
200 °It
t
o30 l
I
I
I
I
I
I
2
4
6
8
I0
12
It+
MgO (wt. %)
mgO
16
(wt.%)
o31
~
e2§ o18
e2++
19
022 11
,o%+:,o,,+,.,0,0+,
.,J
ot
v
0
4 e~
4
o28
~2
o=
2
029
0|
%
O29 22 20] 27 o30 I5
o+.
0
Z •
2
~25
+28
~
6
8
+0
MgO
i
t2
I+
16
2
MgO ( w t . % )
(wt. %)
OZa
70
126
I0
ols
032
016 014
65 o~°
o 23 027
8
,j
,..,+
%
0
oJ1
6
,,o ~oo,,",;"
119
11 120~01
U~?
e16
~0
o+2 02
2~i0Of?
35 30
os~s
2+oi
2~2~o 22 +2s 7 6o°
2s..~g ~o o$ Z~eoI o" °~ 011
g o o3
030 012 o13 029
=tl
v
oa
017
Ol
2
4
6 MgO
I
12
( wt.% )
14
16
,
.
2
8
.
.
10
.
12
}&
McjO ( w t . % )
16
Miocene-Pliocene Hydrothermal Ore Deposits
43
A
I
0
! r~
E
D
Z
10
II
~2
)3
I~
15
m6
J7
Fe 2 0 3 Content (wt.%) of Epidote Figure 1.33. Frequency (number of analyses) histogram for Fe203 (wt%) of epidote from the Kuroko basalt. A: epidote coexisting with albite, B: epidote coexisting with chlorite, C: epidote coexisting with pyrite, D: epidote coexisting with hematite and calcite (Shikazono et al., 1995).
where aAg2S and aAg denote the activity of Ag2S in argentite and the activity of A g in electrum, respectively and fs2 is sulfur fugacity. Barton and T o u l m i n ( t 9 6 4 ) have derived a relationship b e t w e e n fs2, temperature, and the A g c o n t e n t of e l e c t r u m in e q u i l i b r i u m with argentite, u s i n g the equation of W h i t e et al. (1957) for the chemical potential of A g in electrum in c o m b i n a t i o n with the equation
Figure 1.32. The relationship between MgO concentration and other major constituents. Solid circle represents the sample of relatively fresh rock which contains original clinopyroxene. A: H20(+) vs. MgO. B: CaO vs. MgO. C: Na20 vs. MgO. D: K20 vs. MgO. E: SiO2 vs. MgO. F: FeO vs. MgO. G: Fe203 vs. MgO. H: EFe vs. MgO. I: A1203 vs. MgO. (Shikazono et al., 1995).
Chapter 1
44 15
A m J
P;H
10
m
t-
<
"6
$ r~
5
E
I . . . . . .
E-I . . . .
I' I
ill
I11 z.J . l . l . l
0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 MgO/FeO (in wt.%) of Chlorite 15
B
>. ¢,t)
<
_c r~
E z
C
5
i
0
•
I¢IMI~IC c l c l c l IClClClC c l C l c l clclcl IclClClCfc c l c l c l ~ 1 , • .I¢IC1¢1~l~ c1¢1cl
0.2 0.4 0.6 0.8
. ,
1.0 1.2 1.4 1.6 1.8 2.0
MgO/FeO (in wt.%) of Chlorite Figure 1.34. Frequency histogram for MgO/FeO ratios (in wt%) of chlorite from the basalt studied (A) and MORB (B). Data sources are: Shikazono and Kawahata (1987), Humphris and Thompson (1978) (M: Mid-Atlantic Ridge) and Kawahata (1984) (C: Costa Rica Rift, Galapagos Spreading Centre). The data on chlorite from MORB are taken from typical metabasalt and not from quartz-chlorite breccia and veins which formed in a hydrothermaI upflow zone (Shikazono et al., 1987).
of Richardson and Jeffes (1952) for the standard free energy change of reaction, 4Ag (native silver) + $2 ----2Ag2 S (argentite)
(1-4)
The equation derived by them can be expressed as, log f s 2 = (1/4.576T) {- 4 1 9 8 0 + 16.52T - 18.296T log NAg
+ (4(1 -- NAg)2[5650 -- 1600(1 -- NAg) -- 1.375T]) }
(1-5)
where NAg and T denote mole fraction of Ag in electrum and absolute temperature (Kelvins), respectively.
Miocene-Pliocene Hydrothermal Ore Deposits
45
Bedded orebody L-40
~e
,
BSO ,
::.,
:t ,
eoe
L-55
,,
YSO
,
BSO
'
YSO
'
|
i•
(
¢
ee eee eeeeeee
I
I
tee e.e oe ceel.¶•
I
1
i!L
Stockwork orebody L-70
BSO
'
S,
• l eeoe e •e~eee
L- 85
.
BSO ,
YSO
'
|
I
I
I
°Oeeee~eee
I|,;
,o,~o*,e
e
I
t
:|
eee oeoe ole, •lee OQII
k-100
BSO ,
,
,
eel•o••
:;:"F:
....
,
,,"
.
YSO
'
I
i ..e
"e g :t eF,:" b eeee te.eyee.
!
200 ° 300 ° Temperature °C
Figure 1.35. Summarized results of homogenization temperature determination in quartz from Uwamuki No. 4 Orebody shown for Kuroko-type (BSO) and Oko-type (YSO), and siliceous ores and for each level (Marutani and Takenouchi, 1978).
Barton and Toulmin (1964) and Barton (1980) have derived the correction to equation (1)-(5) which is necessary due to the solubility o f Au in argentite as a function o f temperature and electrum composition. Based on this equation and NAg, we can place a limit on fs2 and temperature (Fig. 1.38). This application to ore fluids responsible for Kuroko deposits has been done by Sat• (1969), Kajiwara (1970b) and Shikazono and Shimizu (1988a). B o r n i t e - c h a l c o p y r i t e - p y r i t e assemblage also defines i s z - t e m p e r a t u r e region. Combining the FeS content of sphalerite coexisting with b o r n i t e - c h a l c o p y r i t e - p y r i t e , f s z temperature can be determined (Kouda, 1977).
Chapter 1
46 I
g
Oe
•
i
.
.
.
|
I
.
.
.
.
!
P r i m a r y fluid inclusions JU79SK19
S e c o n d a r y fluid inclusions
0
,.-,.
.
6
o Z
5
O" ¢J
o
(b,
4
sw
1I
e
~2
o?
D
i
Low T
b.
0
Secondary Inclusions (16)
s
&
|
.
.
.
.
• o @
@
. . . .
!
150
. . . .
I
200
. . . .
I
250
Temperature
eQ
300
(°C)
Figure 1.36. Homogenization temperature and salinity of inclusion fluids (Pisutha-Arnond and Ohmoto, 1983)•
•
"
I
Pure Wot =~t-
I I
tso-
LIQUID
//,~'/
-
,,,
=
/
loo
- ,sod
/ //
~'
-.-I=o~o
.
,
I
o - . I.'~
c,O~
/
/
Depth of Seowoter
J
/
/
(meters)
. /
,~
/
,.;.'~1 /
VAPOR
'*
A
Fluid I n c l u s i o n F i l l i n g T e m p s .
100
;~00
300
Temperoture ( % )
I
400
350"¢
Figure 1.37. Estimation of minimum depth of seawater at the time of Kuroko mineralization. P-T diagram of NaC1-H20-CO2 is from Drummond (I981) (Ohmoto et al., 1983).
Miocene-Pliocene Hydrothermal Ore Deposits I
I
|
I
t
47 *
i
!
I
,
-5
-10
O4
09 "0'~
-15
0
--I
-20
-25
I
100
0
!
200
300
Temperature °C Figure 1.38. The formative temperature and f s 2 of the black ore from the Shakanai mine. The arrow mark shows an assumed trend of deposition from earlier to later stage (Kajiwara, 1970b). py: pyrite, bn: bornite, cp: chalcopyrite.
Oxygen fugacity (foz). The relationship betweeen
is 2
and
/O 2 can
be derived from the
following reaction. H z S + 1 / 2 0 2 = H 2 0 + 1/2S2
(1-6)
Equilibrium constant for this reaction (K1-6) is expressed as, 1/2 1/2
K16 = (a.2o/
2)/(aH2s/o2
)
(1-7)
where a = activity. Assuming aH20 = 1 and FHzS (activity coefficient of HzS) = 1, we obtain, log fo2 = log fs2 -- 2 logmHzS -- 2log KI-6
(1-8)
where m = molality. Using this relation, we can estimate fo2, mH2S and fs2 at a given temperature. The temperature and f s 2 c a n be estimated from fluid inclusion homogenization temperatures, and chemical compositions of sphalerite and electrum coexisting with argentite and pyrite. H2S concentration of hydrothermat solution is thought to be controlled by chlorite-pyrite assemblage. It is deduced from the H2S concentration of present-day submarine hydrothermal solution which is 1-10 m m o l / k g H 2 0 (Gamo, 1995). Using estimated fs2, temperature, and mH2S, we can place a limit of f Q .
Carbon dioxidefugacity fc02. CO2 fugacity (fc02) of ore fluids is estimated based on CO2 concentration of fluid inclusions analyzed. By using equilibrium constant of the reaction, CO2(g) + H 2 0 ----H2CO3, and assuming aH20 to he unity, fco2 can be estimated.
Chapter 1
48 ~,t,l,~,,l~[[,I,t[,I,,,tl,,~,l
~,~l,t~[ix,,~lt,~,l~Z,l,~,l
Dolomite+K-mica+~.
0
Magnesite+Kaolinite+Quartz <:3
o
2
0
~--2 0 -r
t~
-4
o
_o =
Caleite+Mg-chlorite +K-feldspar
-6 -8
,,,I
0
50
....
I,,,,I,,,,I,,,jll,,z
100
150
200 Temperature( ~ )
Mg-chlorite "6
250
300
licit
0
50
,lllillLl
100
,I;lll,lllwl~w
150
200
Temperature (*C)
250
300
Figure 1.39. Relationship between fco2, aH2co3 and temperature for silicate--carbonate equilibria. Thermochemical data used for the calculations are taken from Helgeson (1969) (Shikazonoet al., 1998). Magnesite, dolomite and calcite occur in hydrothermally altered rocks near Kuroko orebody. The following equations are used to constrain f c Q values of hydrothermal solutions (Shikazono et al., 1998) (Fig. 1.39). 5CaMg(CO3)2 + (dolomite) =
KA13Si3OIo(OH)2 + 3H20 + 3SIO2 (K-mica) (quartz)
MgsA12Si3OIo(OH)8 + KAISi308 + 5CACO3 + 5CO2 (Mg-chlorite) ( K - f e l d s p a r ) (calcite)
(1-9)
5MgCO3 -t- A12Si2Os(OH)4 -t- 7H20 + SiO2 (magnesite) (kaolinite) (quartz) =
MgsA12Si3Olo(OH)8 + 5H2CO3 (Mg-chlorite)
(1-10)
Since temperature of formation of carbonates can be estimated from homogenization temperature of fluid inclusions in carbonates, we can place a limit of CO2 from the above equilibrium relationships. The estimated CO2 range is 1-0.01 mol/kgH20. 1.3.3.3. Chemical compositions o f ore fluids Assuming chemical equilibrium between Kuroko-forming minerals and Kuroko ore fluids, the chemical composition of Kuroko ore fluids can be estimated, using thermochemical data. Calculations for the purposes of estimating the chemistry of Kuroko
Miocene-Pliocene HydrothermalOre Deposits
49
ore fluids have been carried out by several investigators (Sato, 1969; Kajiwara, 1970b; Shikazono, 1976; Ohmoto et al., 1983). For example, if sphalerite is in equilibrium with aqueous solution, the chemical equilibrium for the following reaction can be derived. ZnS = Zn 2+ + S 2-
(1-11)
Assuming that activity of ZnS is unity, solubility product of ZnS (KI-ll) is expressed as, K i-11 = a z n 2 + a s 2 -
( 1-12)
Activity of S 2- is related to total dissolved sulfur concentration (~S) and pH in the region in which H2S is predominant among aqueous sulfur species. S 2-
+ 2 H + = H2S
(1-13)
The equilibrium relation for this reaction is expressed as,
KI_13=aH2s/(a2+as 2 )
(1-14)
Combining this relation with (1-12), and assuming aH2s = NS, we obtain, 2 K 1-14)= (azn2+]ES)/(a2+Kl-13) Kl-ll=(azn2+au2s)/(aH+
(1-15)
In the Kuroko ore fluids with 1 molal C1 concentration predominant dissolved Zn species are thought to be zinc chloro complexes such as ZnC12. For simplicity, assuming that ZnC12 is predominant among dissolved Zn species (~2Zn = mzncl2=aznc12/YZnCl2, where YZnCl2 = activity coefficient of ZnC12 and it is assumed to be 1, and ~2Zn is total dissolved Zn concentration), the relation between azn2+ and 2 Z n can be derived from the following chemical reaction, ZnC12 = Zn 2+ + 2C1-
(1-16)
The equilibrium constant for (1-16) is, Kl-16 = (KI-11KI-13a2H+/~S)(a21 /NZn)
(1-17)
Therefore, we obtain, log NZn = log K - 2pH + 2 log mcl + 2 log YcI
--
log NS
(1-18)
where K = (K1-I1Kl-~3)/K1-16 and Yc1- = activity coefficient of C I - . Thus, NZn can be calculated if the values of K1-11, Kl-13, KI-16, mci-, Yct- and NS are available. NS in ore fluids is generally in a range of 10-2-10 -3 mol/kg H 2 0 based on IES in present-day geothermal waters and fluid inclusion analytical data (Shikazono, 1972a). YcI- is represented as a function of ionic strength and temperature. Ionic strength is related to salinity which can be approximated as C I - concentration. C1- concentration can be estimated from fluid inclusion study. Concentrations of other metal elements can be also estimated based on the procedure similar to that mentioned above.
Chapter 1
50
An example of the calculated results on the chemical compositions of Kuroko ore fluids is given in Tables 1.5 and 1.6. Because of uncertainties of equilibrium constants, NS, pH, temperature, fo2 and other parameters (activity coefficient, ionic strength, activity of water, pressure), the estimated values of concentrations may have uncertainties of-t-1 in logarithmic unit. However, it can be concluded from the thermochemical calculations and fluid inclusion data that the Kuroko ore fluids have the following chemical features. (1) Slightly acidic. (2) C1- concentration is similar to or higher than that of seawater. (3) Mg concentration is very low, compared with that of seawater. (4) The concentrations of base-metal elements and Ba are considerably higher than those of seawater. Several workers have intended to estimate the chemical compositions of Kuroko ore fluids based on the chemical equilibrium model (Sato, 1973; Kajiwara, 1973; Ichikuni, 1975; Shikazono, 1976; Ohmoto et al., 1983) and computer simulation of the changes in mineralogy and chemical composition of hydrothermal solution during seawater-rock interaction. Although the calculated results (Tables 1.5 and 1.6) are different, they all show that the Kuroko ore fluids have the chemical features (1)-(4) mentioned above. As will be discussed later, the experiments (Hajash, 1975; Mottl and Holland, 1978) and theoretical studies on seawater-rock interaction (Wolery, 1978; Reed, 1983) indicate that the Kuroko ore fluids characterized by (1)-(4) above are formed by seawater-crustal rock interaction at elevated temperatures.
TABLE 1.5 Chemical composition of Kuroko ore solution, estimated based on the solubility data of Kuroko forming minerals (Shikazono, 1985c)
pH Cl Na K Ca Mg Fe SO4
H2S Cu Zn Pb Au Ba Sr T (°C) fo2 (atom)
Kuroko ore solution
Seawater
4.5 :t_0.5 33000 ppm 12500 ppm 4000 ppm >500 ppm 1-10 ppm 300 ppm 10-4 m 5 × 10 3 m 5 ppm 20 ppm 20 ppm 10 3 ppm 30 ppm 1-5 ppm 250°C ~ 50°C
7.89 18800 ppm 10760 ppm 399 ppm 412 ppm I294 ppm 2 x 10 3 ppm 2712 ppm
10 - 3 5 ± I a t o m
0.3 x 10 . 3 ppm 3 x 10 3 ppm 0.03 x 10.3 ppm 0.005 x 10 ? ppm 0.33 x 10 3 ppm 7.9 ppm
Miocene-Pliocene Hydrothermal Ore Deposits
51
TABLE 1.6 Concentrations of sulfur and metal species in the Kuroko ore-forming fluids of pH = 4.5 (Ohmoto et al., 1983) Species
ESO 2-
H2S £Fe ECu IBBa li3Zn EPb EAg H4SiO4
Maximum actual, or minimum value
Concentration (Iog molality)
200°C
250°C
300°C
350°C
Minimum Maximum Maximum Actual Actual Max mum Maximum Minimum Maximum Minimum Maximum Minimum Maximum Minimum Actual
-3.5 -2.0 -2.6 -4.3±0.6 -4.6±1.1 -5.7±0.2 -4.1 -3.0 -3.5±0.6 -4.5±0.6 -4.8±0.6 -5.8±0.6 -5.9±0.3 -6.9±0.3 -2.4
-3.5 -2.0 -2.4 -3.3±0.7 -4.4±1.1 -5.4±0.2 -3.9 -4.7 -3.0i0.7 -4.7±0.7 -4.3±0.7 -6.0~0.7 -4.7±0.4 -6.4i0.4 -2.2
-3.5 -2.0 -2.0 -2.4±0.7 -5.2±1.2 -5.2±0.2 -3.5 -4.2 -2.4±0.7 -4.4±0.7 -3.6±0.7 -5.6±0.7 -4.0±0.4 -6.0±0.4 -2.0
-3.5 -1.5 -0.6 -1.6±0.9 -2.6±1.3 -4.1±0.3 -4.1 -5.4 -2.2±0.9 -5.2±0.9 -2.5±0.9 -5.5±0.9 -2.7±0.5 -5.7±0.5 -1.9
1.3.3.4. Stable isotopes 6D and 6180. 3D and 3180 of the Kuroko ore fluids were estimated based on analyses of fluid inclusions, Kuroko-forming minerals and hydrothermal alteration minerals (e.g., Pisutha-Arnond and Ohmoto, 1983). Estimated 8D and 8180 of Kuroko ore fluids are plotted on 3D-3180 diagram (Fig. 1.40).
40
,
,
.
,
,
,
,
o
,
Kuroke Ore-forming Fl@ds
2O
I
SW ! ~z I Periods ii1,111& IV
Periods I%V
0
B
I I I
~D -20 (%o) -40
- ~-.
~
ndary i n_s J
-60
"Magmatic"FLuid] ' ~
-80 -100 -12
I
-1o
-6
!
-4
I
I
-2 0 6180 (%4
|
I
I
i
2
4
6
8
10
Figure 1.40. Summary of 8D and 8180 values of the Kuroko ore-forming fluids and of low-temperature secondary inclusions (Ohmoto et al., 1983). SW: seawater.
Chapter I
52
They are -30%o to -3%~ for 5D and -2%o to +5%o for ~180. The origin of Kuroko ore fluids based on these data is discussed in section 1.3.4. (]34S ofsulj'qd¢$. A large number of ~348 data on sulfides are available. Figure 1.41 shows the summary of sulfur isotopic compositions of sulfides (chalcopyrite, sphalerite, galena, pyrite) (Shikazono, 1987b). Because pyrite exceeds 80% of total amount of sulfide sulfur in the deposit it is reasonable to assume that 534S of pyrite represents ~34S value of sulfides in Kuroko deposits. ~34S values of sulfide minerals from Kuroko deposits vary widely in a range from -6%0 to +9%o. Pyrite with low ~34S value (-6%0 to 0%o) is
<
GALENA
"5
=E Z
1 -6-5-4-3-2
-
0
1
2
3
4
5
6
,<
7
8
9
10
~3~S(O/oo)
CHALCOPYRITE
"6
Z1 -6
-5
-4
-3
-2
-1
-0
1
2
3
4
5
6
7
8
9
10
~34S(% o)
SPHALERITE
"6
=E Z
1
-6-5
4-3-2-1
-0
1
2
3
4
s
s
~
~
1'0
8~4S(O/c~)
;/ 1
7
PYRITE
,-= ICI E-I E-] -6
-5
-4
-3
~ -2
-1
. -0
. 1
. 2
. 3
4
. 5
6
7
8
9
1'0
~3~S(O/c~)
Figure 1.41. Sulfur isotopic compositions of sulfide minerals from Kuroko deposits (Shikazono, 1987b).
53
Miocene-Pliocene Hydrothermal Ore Deposits
s3~s(%,o) -6 1
mudstones end turfs (ht-qtz zone) .(bar-rich zone) --
•
blaok-ores
-5 1
. i.i . . . . . . . . .
o I
+2 i
f
~= _
•--.
'~ I
+4 i
t
+6 I
+8
t
1
t
1
+10
+22
i
I
__-~_-2-=x_. . . . .
-
+ 24 f
i
~,,~ --
eo
•
bar
,,~ ,.~ ,."
r' ~ j
#
/
yellow-ores pyrite-ores veins and disseminated
~ ]
cp
~
,,, I,,PY
Miooene sea
ores
Figure 1.42. Sulfur isotopic variation and the vertical zonation of ores in the Shakanai No. 1 deposit (Kajiwara, 1971).
framboidal type (Komuro and Sasaki, 1985). The origin of framboidal pyrite is not well understood. However, it seems likely that this type of pyrite formed biogenetically and is not of hydrothermal origin. Most ~34S values of hydrothermal pyrite having euhedral morphology range from +2%0 to +7%0. The approximate ranges of variation of ~34S values of other sulfides are: chalcopyrite; +2%0 to +7%0; galena: 0%0 to +4%0; sphalerite: +2%0 to +7%0 (Sasaki, 1974). Sasaki and Kajiwara (1971) estimated average ~34S value of sulfide minerals to be +4.6%0 based on the measurements of mill concentrates from representative Kuroko deposits. In individual deposits, ~34S of sulfides generally increases stratigraphically upwards (Fig. 1.42). (Kajiwara, 1971). Based on the sulfur isotope evidence, Kajiwara (1971) deduced that the ore solutions underwent a progressive cooling and oxidation due to mixing with seawater. Sulfur isotopic data of separated pyrite as the commonest sulfide mineral (Kajiwara, 1971; Kajiwara and Date, 1971) show different ~348 values for the three sub-types of Horikoshi and Shikazono (1978). The ~34S values of pyrite in the C sub-type deposits are higher than the ~34S values of pyrite from the Y and B sub-types. The ~34S values of pyrite from the Y sub-type seem to be slightly higher than those from the B sub-type. Kajiwara and Date (1971) are of a different opinion: the 334S values from the Kosaka district are higher than those in the Hanaoka district, because all sulfur isotopic data from the C sub-type were obtained in the Kosaka district. The sulfur isotopic data on the obtained Uwamuki deposits of the B sub-type in the Hanaoka district indicate systematic decrease in ~348 passing from the yellow ore (+7%0) to the black siliceous ore (+5%o) (Bryndzia et al., 1983). Kajiwara and Date's data (1971) include three ~348 values of pyrite in the Doyashiki deposit of C sub-type in the Hanaoka district. The main Doyashiki
Chapter 1
54 5/+.9
5'
+40
SD
YU
+
AK +35
~.~.-30
TN06
CO
NK
"KS, f t '
•
rN •.~e
I"IZ HT ll.,rr
''I +"H""",
V
+20 1<13 IN
OG
.15 *5
.10
+15
*20
so(sMow) Figure 1.43. Relationship between 3t80 and ~34S values of barite from some Kuroko and vein-type deposits. Abbreviations are: HN, Hirano; NK, Nagaki; AK, Akakura; YU, Yunosawa; SD, Sado; KY, Katsuyama; TN, Teine; OS, Osarizawa; MZ, Mitsuzawa; HT, Hata; OE, Oe; AI, Akaishi; KN, Kohinata; MT, Miyatamata; KZ, Karuizawa; FN, Funauchi; KO, Koyama; IK, Inakuraishi; and OG, Ogoya. "K.S.R." indicates "Karoko sulfate region" (Watanabe and Sakai, 1983).
deposit was, however, mostly mined out when their samples were collected in about 1969. Furthermore, the mine calls the main Doyashiki of the C sub-type and the satellite deposits of Y sub-type together with the Doyashiki deposits (Takahashi and Suga, 1974).
and 3180 of sulfates. 834S and 8180 of sulfates (anhydrite, gypsum, barite) of Kuroko deposits were reported by Sakai et al. (1970) who indicated that 834S and 8180 values of sulfates are very close to but slightly higher than Miocene seawater sulfate value (834S = +20%o to +21%o; 8180 = 0%o). After Sakai's pioneering work, Watanabe and Sakai (1983) and Kusakabe and Chiba (1983) analyzed large amounts of barite, gypsum and anhydrite for 834S and 8180 and confirmed the conclusion drawn by Sakai et al. (1970) (Fig. 1.43). t]34S
1.3.3.5. Radiogenic isotopes Lead isotopes. Sato and Sasaki (1973) concluded on the basis of a remarkable narrow range in lead isotopic composition of Kuroko ores that lead of Kuroko ore came from deep-seated source which originated from subducting pelagic sediments.
Miocene-Pliocene Hydrothermal Ore Deposits o
38.8"
o
/"~--'~Hokuroku ,,4" j ores / ~ .~
central
38.6.
2ospb 2°4Pb 38.4,
55
Honshu.,,/~'.
o
i. ', I o i
•
382"
.p">.%~ j Z
38.0.
...~/
' / I
•
I• o I"iI I /" 0 /" / / , ~ i . / / I- "II /
~
I
[northern /\Honshu
/"
_.
o Kuroko • Epigenetic
t
rowth curve
207pb 15.6- !
2o4pb
~(7 o
15.5-
~
"?.'"~ "--£'o • /
•
/ ~ Honshu
15.4- centrcl| Honshu
18~
'
1~4
'
2o6pb/2O4pb
ld.6
Figure 1.44. Isotopic composition of lead in black ore (open symbols) and in yellow ore (closed symbols) in the Hokuroku district. The isotopic fields for black ore from the Fukazawa, Shakanai, and Kosaka deposits are outlined (Fehn et al., 1983).
Lead isotopic data on Kuroko deposits, vein-type deposits in Honshu and volcanic rocks are summarized and plotted in Fig. 1.44 (Fehn et al., 1983)• Although lead isotopic compositions of Kuroko ores occupy a narrow isotopic range, within a given ore deposit, black ore has a uniform isotopic composition but is significantly higher in radiogenic lead than yellow ore (Fig. 1.44, Table 1.7; Fehn et al. 1983). On the basis of the lead isotopic distribution, Fehn et al. (1983) concluded that a major part of the lead in Kuroko deposits was derived from igneous rocks, probably volcanic rocks with a significant contribution coming from the underlying preNishikurozawa formations and the yellow ore seems to have a greater igneous rock lead component than does the black ore.
Strontium isotopes. Strontium isotopic compositions (87Sr/86Sr) of anhydrite, gypsum and barite from Kuroko deposits are summarized in Fig. 1.45 (Farrell et al., 1978; Honma and Shuto, 1979; Farrell and Holland, 1983; Yoneda et al., 1993; Yoneda and Shirahata, 1995). 87Sr/86Sr values of anhydrite and gypsum are slightly lower than that of seawater, suggesting that most of the strontium was derived from seawater, but a small amount of
Chapter 1
56 TABLE 1.7 Isotopic composition of leads from Kuroko deposits (Fehn et at., 1983) Sample number
Field number
2°6pb/2°4pb
2°Tpb/2°4pb
2°8pb/2°4pb
W 1112132 K- 105 W 1112125
18.489 18.463 18.428
15.603 15.589 15.567
38.678 38.623 38.561 a
JU 78UM64 JU 78UM67
18.458 18.475
15.584 15.594
38.577 b 38.619 b
76-12-2-3 JU 7SUMI25 JU 78UM 129 CF 082713 CF 111163B CF 11 1163Y
18.469 18.464 18.470 18.473 18.491 18.434
15.594 15.5s9 15.594 15.589 15.604 i5.560
38.618 b 38.592 b 38.618 b 38.622 38.665 38.508 c
CF CF CF CF CF CF CF CF CF CF CF CF
18.550 18.580 18.568 18.565 18.566 18.558 18.560 18.575 18.561 18.546 18.542 18.518
15.621 15.622 I5.607 15.617 15.607 15.615 15.595 15.604 15.618 15.596 15.589 15.592
38.686 38.736 38.729 38.659 38.702 38.698 38.665 38.734 38.709 38.637 c 38.613 c 38.620 a
CF7 11282 CF 11286 S 11
18.527 t8.536 18.479
15.621 13.620 15.583
38.707 38.701 38.584 a
S 30
18.527
I5.617
38.624
S 35
18.539
15.598
38.683
MT9
18.451
15.567
38.517
M 8
18.448
15.585
38.378
Kosaka area
Uchinotai East 1 2 3
Uwamuki 2 4 5
Uwamuki 4 6 7 8 9 10 11 Fukazawa area
Tsunokakezawa 12 13 14 15 16 17 18 19 20 2I 22 23 Matsurnine-Sbakanai
11197 34511 l 11922 I 11810 34516 110197 082610B l 11913 111813 I1184 682610Y 111920
area
Shakanai 4 24 25 26
Shakanai 8 27
Shakanai 11 28
Matsumine 29
Matsuki 30
If not stated otherwise all samples were measured on galena from black ore. a Galena taken from yellow ore; b From Sato et al. (1981); c Yellow ore sample.
Miocene-Pliocene Hydrothermal Ore Deposits
57
TABLE 1.7 (continued) Sample number Other
Field number
2°6pb/2°4pb
2°7pb/2°4pb
2°spb/2°4pb
HA 41
18.521
15.587
38.623
CF 11233
18.525
15.594
38.626
FU 17
18.556
15.603
38.662
deposits
Hanawa
31 Omaki
32 Furutobe
33
T
~10
[ ] Gypsum []
c ,< "6
Anhydrite
5
E "1 Z
1
,
.
.m..N.I 0.7070
N N N ~ ~
N N , 0.7080
N N N
0.7090 87Sr/86Sr
Figure 1.45. Variation of the 87Sr/86Sr ratio of anhydrite and gypsum from Kuroko deposits (Farrell and HoIiand, 1983; Shikazono et aI., 1983).
strontium was from igneous rocks. 87Sr/86Sr ratios of barite are in a range from 0.706 to 0.708, suggesting smaller contribution of seawater strontium than anhydrite and gypsum. 87Sr/86Sr of barite from small Kuroko deposits (Iwami in San-in and Minamishiraoi in Hokkaido) are lower (Yoneda, et al., 1993; Honma and Shuto, 1979; Farrell and Holland, 1983) than that from large Kuroko deposits in Hokuroku district. This indicates that a large seawater circulation did not occur in a small Kuroko mine area and that seawater circulation is important for the formation of Kuroko deposits.
Rare earth elements (REE). Analytical results of REE contents of hydrothermally altered volcanic rocks in Kuroko mine area and Kuroko ores are summarized as follows (Shikazono, 1999a) (Fig. 1.46). (1) Positive Eu anomaly is observed for barite, Kuroko ores, ferruginous chert (tetsusekiei), and hydrothermally altered basaltic and dacitic rocks overlying the Kuroko ores. (2) Negative Eu anomaly is observed for hydrothermally altered dacite underlying the Kuroko ores and anhydrite in the dacitic tuff breccia.
Chapter I
58
Hydrothermally altered il dacite and anhydrite underlying the i Kuroko ores E lO0
Barite, Kuroko
i
ore ant
Hydrothermally altere basalt overlying the Kuroko ores
err i
i i
::
::
Ns.s~
i
NS-~S
::
!
i
~
sa2
"-- NS-33
NS-36
NS-39
NS.2
i La Ce
• Nd
i SmEu
i Tb
i i Yb L~
i La Ce
i SntEu
i Tb
i i Yb Lu
I La Ce
i Nd
i SmEu
i Tb
t t Yb Lu
Figure 1.46. REE patterns of the altered volcanogenic rocks and Kuroko ores. Data sources: Shikazono (1999a). (A) Hydrothermally altered dacite and anhydrite underlying the Kuroko ores. (B) Barite, Kuroko ore and ferruginous chert. (C) Hydrothermally altered basalt overlying the Kuroko ores (Shikazono, 1999a).
(3) N e g a t i v e Ce a n o m a l y is o b s e r v e d for h y d r o t h e r m a l l y altered chlorite-rich basalt o v e r l y i n g the K u r o k o ores. (4) N e g a t i v e Ce a n o m a l y and positive Eu a n o m a l y are o b s e r v e d for epidote-rich altered basalt near the orebody. (5) L i g h t rare earth e n r i c h m e n t is distinct and R E E contents are relatively high for the ferruginous chert. The R E E pattern for fresh v o l c a n i c rocks in the Kuroko m i n e area studied by Dud~is et al. (1983) is shown in Fig. 1.47 which shows no negative Ce and no positive Eu a n o m a l i e s and L R E E (Light Rare Earth E l e m e n t ) are not enriched c o m p a r e d with H R E E
1000 Sample / Chondrite
1000~ ~_S a m p l e / C h o n d r i t e
a
Basaltic Rocks
b
~
Acidic Rocks
I lOO
~
,,,..
I0
1 La
!...... _ " : 2 _ _ _ -
10
Ce
Sm Eu
Tb
YbLu
La C e
Pr
Nd
S m E u G d T b D y Ho E r
Yb
Figure 1.47. REE patterns of the fresh volcanic rocks in the Kuroko mine area. Data source: Dud~s et al. (1983). (a) basalt; (b) acidic rocks (Shikazono, I999a).
Miocene-Pliocene Hydrothermal Ore Deposits
59
(Heavy Rare Earth Element). Therefore, it is considered that negative Ce and positive Eu anomalies in hydrothermally altered volcanic rocks, Kuroko ores, and ferruginous chert and LREE enrichment in the Kuroko ores have been caused by hydrothermal alteration and precipitations of minerals from hydrothermal solution responsible for sulfides-sulfate (barite) mineralization. Negative Eu anomaly is found for anhydrite sample and altered dacite underlying the Kuroko ores. One of the possible explanations for this negative anomaly is selective leaching of Eu by circulating hydrothermal solution of seawater origin. Sverjensky (1984) has shown from thermochemical calculations that Eu 2+ is more abundant than Eu 3+ under the reduced environment at high temperatures. Thus, it is considered that Eu is leached more efficiently from the rocks compared with the other rare earth elements. Alderton et al. (1980) have shown that Eu in the granitic rocks in southwest England is depleted due to the sericitization of feldspar-bearing assemblage. Negative Eu anomaly has been reported on highly silicified volcanic rocks around the volcanic rocks around the volcanogenic polymetallic massive sulfide deposit at Que River, Tasmania (Whitford et al., 1988). Thus, it is plausible that negative Eu anomaly of dacitic rocks was caused by the sericitization. This negative Eu anomaly indicates that the alteration minerals in the volcanic rocks underlying the Kuroko ores did not precipitate from ascending hydrothermal solution which interacted at relatively high temperatures (more than 250°C) and reduced condition (Eu2+/Eu 3+ is more than 1) and have positive Eu anomaly. Date et al. (1983) and Green et al. (1983) have shown based on numerous analytical data that Na, Ca and Sr were depleted from footwall dacitic rocks below the Fukazawa Kuroko deposits and this depletion was caused by the addition of K from ascending hydrothermal solution accompanied by the destruction of plagioclase, K-feldspar and volcanic glass and formation of sericite. This evidence supports the above interpretation: The selective leaching of Eu from footwall dacite was caused by the hydrothermal solution. The REE pattern for anhydrite is different from that of seawater, indicating that anhydrite did not precipitate due to the simple heating of seawater that was suggested by Sakai et al. (1970) and Sato (1973). This REE pattern could be explained in terms of the mixing of hydrothermal solution and cold seawater and low degree of seawater/hydrothermal solution mixing ratio (Shikazono et al., 1983). Negative Eu anomaly is also found in the fresh and altered dacitic rocks (Dudfis et al., 1983). Therefore this negative anomaly in anhydrite is also explained in terms of an influence of sericitization of dacite accompanied by the depletion of Eu. Sverjensky (1984) calculated the dependency of Eu2+/Eu 3+ in hydrothermal solution on f Q (oxygen fugacity), pH and temperature. According to his calculations and assuming temperature, pH and fo2 for epidote-stage alteration of basalt and Kuroko ores (Shikazono, 1976), divalent Eu is considered to be dominant in the rocks and hydrothermal solution. Thus, it is reasonable to consider that Eu in the rocks was removed to hydrothermal solution under the relatively reduced condition more easily than the other REE which are all trivalent state in hydrothermal solution. Thus, it is likely that Eu is enriched in epidote-rich altered volcanic rocks. Probably Eu was taken up by the rocks from Eu-enriched hydrothermal solution which was generated by seawater-volcanic rock interaction at relatively low water/rock ratio.
60
Chapter 1
A negative correlation between Mg content and Ca content of hydrothermally altered basalt and dacite from the Kuroko mine area exists. This correlation indicates that Ca in the rocks is removed to fluid by the exchange of Mg in seawater. Eu may behave in the manner similar to Ca during seawater-volcanic rock interaction because of the similarity of their ionic radii. Positive Eu anomaly is observed for hydrothermal solution issuing from the hydrothermal vent on the seawater at East Pacific Rise (Bence, 1983; Michard et al., 1983; Michard and Albarbde, 1986). Guichard et al. (1979) have shown that the continental hydrothermal barites have a positive Eu anomaly, indicating a relatively reduced environment. Graf (1977) has shown that massive sulfide deposits and associated rocks from the Bathurst-Newcastle district, New Brunswick have positive Eu anomalies. These data are compatible with positive Eu anomaly of altered basaltic rocks, ferruginous chert and Kuroko ores in Kuroko mine area having positive Eu anomaly and strongly support that Eu is present as divalent state in hydrothermal solution responsible for the hydrothermal alteration and Kuroko mineralization. Hydrothermally altered basalt in Kuroko mine area can be divided into chloriterich and epidote-rich one (Shikazono et al., 1995). Chlorite formed under the higher water/rock ratio and lower temperatures than epidote (Shikazono, 1984; Shikazono and Kawahata, 1987). At the stage of chlorite formation, a large amount of seawater cycled (seawater/basalt ratio, 40 by mass) (Shikazono et al., 1995). It is considered that Ce4+/Ce 3+ was high during the interaction of relatively low-temperature seawater and rocks at this stage. Ce is present as trivalent in the original rocks. With the proceeding of the reaction of oxidized fluid (modified seawater) with volcanic rocks, Ce is removed to fluid as Ce 4+ which is stable in oxidized fluid as complexes. Removed Ce 4+ may be incorporated into Mn-hydroxides and oxides (e.g., Goldberg et al., 1963; Courtois and Clauer, 1980). It is commonly observed that the weathered rocks exhibit a negative Ce anomaly (Ludden and Thompson, 1979; Menzies et al., 1979). Although these Mn minerals are not found in the Kuroko mine area, a positive Ce anomaly may be caused by the fixation of Ce 4+ in the altered rocks. The negative anomaly is interpreted in terms of the above consideration. On the contrary, the other rare earth elements (except for Eu) in the rocks are present as trivalent. Thus, it is likely that Ce 4+ is more soluble than the other rare earth elements by the reaction of oxidized seawater with rocks. Thus, the negative Ce anomaly of altered basalt could be interpreted in terms of low-temperature interaction of oxidized and relatively unreacted seawater having a negative Ce anomaly caused by the incorporation of Ce into Mn-hydroxides with rocks at late-stage of submarine hydrothermal activity. The REE data, combined with alteration minerals and concentration of major elements in hydrothermally altered rocks, could be used to reconstruct the structure and evolution of a submarine geothermal system accompanied by Kuroko mineralization (Shikazono, 1999a). At the stage of Kuroko mineralization, evolved reacted seawater enriched in Eu, Ca, and Sr formed at low seawater/rock ratio (ca. 1 by mass) and at relatively reduced condition (Eu2+/Eu 3+ greater than 1). Selective leaching of Eu, Ca and Sr occurred from the dacitic rocks underlying the Kuroko ores. The hydrothermal solution enriched
Miocene-Pliocene Hydrothermal Ore Deposits
61
in Ca, Eu and Sr issued from the discharge zone on the seafloor and sulfides and sulfates were precipitated from such hydrothermal solution due to the mixing with cold seawater. Positive Eu anomaly is strong and REE contents are low for the barite-rich ores, whereas positive Eu anomaly is weak, but REE contents are high for the ferruginous chert ore. This difference suggests that hydrothermal signal (low REE and positive Eu anomaly) was strong and was preserved at the stage of the precipitation of sulfides and barite and was weak and not preserved at the ferruginous chert stage. Probably the high REE concentrations of the ferruginous chert is due to the effect of bottom-seawater mixing causing scavenging of REE in cold seawater by Fe- and Mn-hydroxides and oxides forming the ferruginous chert. This view is consistent with the results obtained by Klinkhammer et al. (1983), Ruhlin and Owen (1986) and Olivarez and Owen (1989) who showed the scavenging of REE in seawater by Fe-oxides and Fe-hydroxides near midoceanic ridges. The effect of bottom-water mixing was discussed also for Archean iron formations and Red Sea metalliferous sediments by Barrett et al. (1988). The REE study indicates that the concentrations of REE, particularly Eu and Ce in altered rocks and ore minerals are useful indicators of oxidation state, intensity of discharging hydrothermal solution and evolutionary stage of submarine hydrothermal activity. 1.3.4. Depositional mechanism and origin of ore fluids
1.3.4.1. Depositional mechanism Some mechanisms of anhydrite deposition in Kuroko deposits. Shikazono et al. (1983) considered the depositional mechanism of anhydrite based on the mode of occurrence, texture, Sr content, nature of the contained fluid inclusions and isotopic composition of Sr, S and O in anhydrite together with the mineralogy of the sekko ore, combined with their experimental study on the patitioning of Sr between coexisting anhydrite and aqueous solution. The following is their discussion on the depositional mechanism of anhydrite. Several mechanisms could be invoked to explain the deposition of anhydrite in Kuroko deposits. These include: (1) recrystallization of gypsum and/or anhydrite of evaporite origin; (2) precipitation of anhydrite due to the cooling of hydrothermal solutions; (3) precipitation of anhydrite due to the boiling of hydrothermal solutions; (4) replacement of calcic minerals, such as feldspars, in volcanic rocks or of calcareous foraminifera in mudstones; (5) simple heating of seawater without interaction with country rocks, either above the seawater-sediment interface (Kajiwara, 1971) or beneath the seawater-sediment interface (Farrell et al., 1978; Farrell and Holland, 1983); (6) heating of seawater accompanied by interaction with country rocks; and (7) mixing of ascending hydrothermal solutions with seawater at a site either above the seawater sediment interface (Sato, 1973) or beneath the seawater-sediment interface. Several of these mechanisms can be ruled out based on the geologic environment of the Hokuroku district, because the geologic environment of the Hokuroku area during the Miocene was quite different from that of areas of evaporite formation. The precipitation of anhydrite from hydrothermal solutions has been studied extensively by various workers (e.g., Marshall et al., 1964a,b). The salinity of the inclu-
62
Chapter 1
sion fluids is less than ca. 5 wt% (~1 mol/kg H20) (e.g., Marutani and Takenouchi, 1978; Pisutha-Arnond and Ohmoto, 1983). Simple cooling of hydrothermal solutions is therefore virtually ruled out as a mechanism of anhydrite deposition. This also holds true for the proposal that anhydrite precipitation was due to boiling of the hydrothermal solutions. Studies of fluid inclusions in minerals from Kuroko deposits have yet to produce evidence of boiling in the inclusion fluids (e.g., Marutani and Takenouchi, 1978). Kumita et al. (1980) have investigated the distribution of foraminifera in the M1 and M2 mudstones that overlie the sulfide horizons of the Shakanai mines and have found that the relative abundance of calcareous foraminifera compared to arenaceous foraminifera decreases toward the sulfide orebody. Higher ratios of calcareous foraminifera to arenaceous foraminifera are found in mudstones stratigraphically higher than the thicker sekko body. This evidence suggests that some of the calcium in the anhydrites may have been derived from calcareous foraminifera in the mudstone. However, anhydrite generally occurs in the uppermost part of the tuff and tuff breccia of the T3 unit, which underlies the M2 and M I mudstones. The quantity of anhydrite in the mudstones is very small. This suggests that most of the calcium in the sekko anhydrites was not derived from the calcareous foraminifera distributed in the mudstones. Foraminifera are not found in the T3 tuff unit, and textures showing the replacement of foraminifera by anhydrite have not been observed. If seawater is simply heated, anhydrite begins to precipitate at approximately 110°C if PH20 is equal to Ptotal. The concentration of Ca 2+, Sr 2+ and SO42- in Miocene seawater was probably similar to that of present-day seawater (Graham et al., 1982). The partition coefficient of Sr between anhydrite and aqueous solution has been measured (Shikazono and Holland, 1983) as has the solubility of anhydrite in seawater (Marshall et al., 1964a,b); the variation in the Sr content of anhydrite with temperature during simple heating of seawater can therefore be calculated, using a value of 0.25 for the partition coefficient, Kdsr, experimentally determined by Shikazono and Holland (1983). Figure 1.48 shows that the Sr content of anhydrite which initially precipitated from seawater at about 110°C, should be about 1500 ppm; the Sr content of the anhydrites from Kuroko deposits is between 200 and 2,000 ppm (Fig. I. 19). The low Sr content of most of these anhydrites is difficult to explain in terms of the simple heating of seawater unless the experimentally determined values of KdSr (Shikazono and Holland, 1983) are not applicable to anhydrite deposition in Kuroko deposits. The relationship between the Sr content and the 87Sr/86Sr ratio of a number of anhydrites from Kuroko deposits is shown in Fig. 1.49. Most of the data points fall close to a trend line along which the Sr content of anhydrites increases with decreasing 87Sr/86Sr. The samples from the Fukazawa mines include sekko and paragenetically late vein anhydrites. The line marked A in Fig. 1.49 represents the position of anhydrites precipitated during the heating of Miocene seawater if KdSr had a value of 0.24 during the process of anhydrite precipitation in Kuroko deposits. The large separation between line A and the analytical data for the Kuroko anhydrites is striking. Miocene seawater was apparently not the only source of Sr in these anhydrites. Unfortunately, the processes by which the trend line of the anhydrite compositions was generated are still impossible
Miocene-Pliocene Hydrothermal Ore Deposits
63
3000
"6 E
~2000 c 0
0 1000
w
60
8
I
I
i
I
i
1O0 120 140 Temperature('C )
l
160
Figure 1.48. Change in the strontium content of anhydrite precipitated during the heating of normal seawater without any seawater-rock interaction (Shikazono et al., 1983).
87Sr/86Sr
0.709
R=0.1 0.2
0.3
0.4
0.5. .
. . .
A
B
C /
0.708 I
0.707
[l ./
lJ !a t it
1!
tl
0.706
0.705
lb!
5 0
i
J
1000
2000 S r c o n t e n t (pprn.)
Figure 1.49. Change of the strontium content and 87Sr/86Sr ratio of Kuroko anhydrite during the deposition and dissolution due to the mixing of hot ascending solution and cold solution (normal seawater) (Shikazono et aI., 1983). R mixing ratio (in weight) = S.W./(S.W.+H.S.) in which S.W. and H.S. are seawater and hydrothermal solution, respectively. Open triangle: Fukazawa deposits, Solid triangle: Hanawa deposits, Open square: Wanibuchi deposits, Solid square: Shakanai deposits. Concentration of Ca 2+, Sr 2+ and SO 2 of H.S. are assumed to be 1,000 pprn, 1 ppm, and 10 -4 mol/kg H20, respectively. Concentrations of Ca 2+, Sr 2+ and SO 2 of S.W. are taken to be 412 ppm, 8 ppm, and 2,712 ppm. Temperatures of H.S. and S.W. are assumed to be 350°C and 5°C (Shikazono et aI., 1983).
64
Chapter 1
to define uniquely. The anhydrites containing the lowest concentration of Sr have an isotopic composition close to that of Miocene seawater. This can be explained in three ways: (1) by a value of KdSr much lower than 0.24 during the precipitation of anhydrite in Kuroko deposits, (2) by the reaction of seawater with country rocks, and (3) by the mixing of seawater with one or more solutions in which the Sr/Ca ratio is much smaller than that of seawater. The first alternative is unlikely but not impossible. Past experience suggests that Kd Sr probably depends on the degree of supersaturation of the solutions with respect to calcite during calcite precipitation (Katz et al., 1972). In the experiments by Shikazono and Holland (1983), the solutions from which anhydrite was deposited were considerably supersaturated. It is therefore possible that the values of KdSr extracted from their experimental data are higher than those which controlled the incorporation of Sr in anhydrite during the formation of Kuroko deposits. Experiments at very low degrees of anhydrite supersaturation are needed to determine whether this is a possible explanation for the low Sr content of some of the Kuroko anhydrites. The reaction of seawater with country rocks is also a possible but unlikely explanation. Tertiary volcanic sediments in the vicinity of Kuroko deposits are altered and tend to have lost both Ca and Sr (Farrell and Holland, 1983). The ratio of Sr loss to Ca loss is roughly equal to the Sr/Ca ratio in seawater. If seawater was the altering medium, its St/Ca ratio was probably not strongly affected by the alteration process. The 87Sr/86Sr ratio would be intermediate between an initial value of 0.7088 and ca. 0.740 the 87Sr/86Sr ratio of unaltered Tertiary volcanics of the Hokuroku basin. It is unlikely, therefore, that this type of alteration can account for the Sr content and for the isotopic composition of Sr in the anhydrites at the upper end of the trend line in Fig. 1.49. On the other hand, mixing of seawater with solutions which have a St/Ca ratio much smaller than that of seawater could have led to the deposition of Kuroko anhydrites. Mixing with solutions containing high concentrations of Ca 2+ and very little Sr 2+ or SO] could lead to the precipitation of anhydrite whose Sr content reflects the low St/Ca ratio of the resulting mixtures. Their isotopic composition would be close to that of seawater. If the solutions that mixed with seawater contained low concentrations of Sr with an isotopic composition of ca. 0.7050, anhydrites precipitated from such mixtures would have tended to lie along a family of curves such as curve B in Fig. 1.49. The position of these curves depends on the concentration of Sr 2+, Ca 2+, and SO 2- in the mixing solution as well as on the isotopic composition of its Sr. It seems unlikely that mixing with a series of such solutions is a reasonable explanation for the trend of the data for Kuroko anhydrites in Fig. 1.49. It seems more likely that this trend was generated by the mixing of solutions from which the low Sr, high 87Sr/86Sr anhydrites were deposited with solutions characterized by a considerably higher Sr/Ca ratio and an SVSr/S6Sr ratio of ca. 0.7070. These solutions could have been the hydrothermal solutions from which the barites and presumably the sulfides in the several Kuroko mine areas were deposited. The 87Sr/S6Sr ratio of barites falls in the range 0.7069 to 07079 (Farrell and Holland, 1983): unfortunately the Sr/Ca ratio of the solutions from which the barites were deposited is poorly constrained. If two solutions containing Sr and Ca are mixed, and if no Sr or Ca containing phase is precipitated, the 87Sr/86Sr ratio of the mixtures is related to their -
-
Miocene-Pliocene Hydrothermal Ore Deposits
65
Sr/Ca ratio by the relationship:
{(87 Sr/86Sr) _ (87 Sr/86Sr)1 }/{(87 Sr/86Sr) _ (87 Sr/86Sr)2 }
,~{(mca2+/msr2+)-(mca2+/msr2+)l}/{(mca2+/msr2+)-(mca2+/msr2+)2
} (1-19)
where m denotes molality. The bar indicates the properties of the mixtures, and the subscripts 1 and 2 indicate those of the two end-member solutions. If the amount of anhydrite precipitated during mixing is sufficiently small so that the Sr/Ca ratio of the solutions is not thereby affected, the Sr content of the anhydrites will be related to the isotopic composition of the contained Sr by curves such as curve C in Fig. 1.49. If the quantity of anhydrite precipitated from the mixture is sufficiently large so that the Sr/Ca ratio of the solutions is affected significantly, the Sr content of the anhydrites of any given 87Sr/S6Sr ratio will be greater than those along the calculated mixing curve. The data for the concentration and the isotopic composition of Sr in the analyzed Kuroko anhydrites are generally consistent with such a mixing model. Unfortunately, the model is not unique, and additional chemical and/or isotopic constraints are needed to test its validity. Ohmoto et al. (1983) and Kusakabe and Chiba (1983) also reached the conclusion that the 8180 vs. 87Sr/86Sr relationship and ~34S vs. temperature relationship of barite from the Fukazawa deposit in the Hokuroku district may be explained by a mixing model with a seawater contribution of less than 20% at temperatures around 200°C. Sato (1973) and Ohmoto et al. (1983) calculated the amounts of sulfides precipitated due to the mixing of ascending hydrothermal solution and cold seawater. Their calculations showed that the calculated ratios of the amounts of minerals precipitated are generally consistent with those in Kuroko ore deposits. At early stage of mineralization, anhydrite formed by the mixing of seawater with hydrothermal solution which interacted in shallow part with volcanic rocks at seawaterdominated condition by decreasing Sr/Ca ratio of seawater but not so changing 87Sr/86Sr. In this stage Mg-chlorite formed from such hydrothermal solution dominantly originated from seawater. After this stage, sulfides, barite, and quartz precipitations took place at the sub-seafloor and on the seafloor by the rapid mixing of hydrothermal solution with cold seawater. Rapid precipitation of minerals occurred from the supersaturated solution under non-equilibrium condition. Equilibrium processes cannot explain the following two points. (1) ~34S values of coexisting sulfate and sulfides, and (2) barite/quartz ratio in orebody. If sulfur isotopic equilibrium between coexisting sulfates and sulfides was attained, using average ~34S values of sulfates and sulfides, +22%o and +5%o, respectively, we could estimate temperature using the equation by Ohmoto and Rye (1979). This temperature seems too high compared with temperature estimated from fluid inclusions and mineral assemblages (section 1.3.3). That means that sulfates and sulfides precipitated under the condition far from equilibrium. If hydrothermal solution in which H2S is dominant aqueous sulfur species and is free mixed with cold seawater in which high amounts of SO]- are contained,
SO]-
Chapter 1
66 6O 5O
• Black ore
40
• Yellow ore
30
• Siliceous ore L)
20
¢=
10
•
00
#~ 10
~., 20
$ 30
=-40
, 50
. 60
70
Quartz Content [wt% ]
Figure 1.50. Quartz and barite content of Kurokoore. but no H2S, it is thermochemically predicted that both quartz and barite precipitate with increasing the cold seawater/hydrothermal solution ratio because solubility of quartz decreases with decreasing of temperature and that of barite decreases with increasing S O ] - concentration, which means decreasing of temperature. However, the barite and quartz contents of Kuroko orebody do not positively correlate with each other (Fig. 1.50).
Precipitation of barite and quartz. Barite and quartz are the most common gangue minerals in the submarine hydrothermal ore deposits such as Kuroko deposits and backarc basin deposits (e.g., Okinawa, Mariana deposits) (Halbach et al., 1989; Shikazono, 1994; Shikazono and Kusakabe, 1999). These minerals are also common in midoceanic ridge deposits. The observations of hydrothermal vents at midoceanic ridges and back-arc basins indicated that the minerals precipitate from the hydrothermal solutions which mix with cold ambient seawater. The mineralogical studies of the chimneys from these areas clarified that metastable phases (e.g., amorphous silica, wurtzite, marcasite, native sulfur) are common in the chimneys. The metastable phases are thought to have formed from the highly supersaturated solution. The degree of supersaturation (or saturation index) for barite in Kuroko deposits is estimated to be less than 100 (Shikazono, 1994). The degree of supersaturation with respect to quartz solubility is high because amorphous silica is precipitating from the solution. The above lines of evidence suggest that the precipitation of minerals in submarine hydrothermal ore deposits on the seafloor is taking place from the fluids with high flow rate at the orifices of the chimney (ca. 1-10 m/s) and with high degree of supersaturation under the non-equilibrium conditions. Gamo (1995) revealed based on the chemical and isotopic compositions of hydrothermal fluids from midocean ridges that the precipitation of minerals and interaction
Miocene-Pliocene Hydrothermal Ore Deposits
67
of fluids and sediments under the seafloor affect the chemistry of fluids discharging from the seafloor. Shikazono et al. (1983) indicated that the anhydrite in Kuroko deposits formed at subseafloor environments. Some applications of the coupled fluid flow-reaction model were carried out to the ore-forming process (e.g., Lichtner and Biino, 1992). However, a few attempts to understand quantitatively the precipitations of minerals from flowing supersaturated fluids in the submarine hydrothermal systems have been done (Wells and Ghiorso, 1991). Wells and Ghiorso (1991) discussed the silica behavior in midoceanic ridge hydrothermal system below the seafloor using a coupled fluid flow-reaction model. The behavior of silica and barite precipitation from the hydrothermal solution which mixes with cold seawater above and below the seafloor based on the thermochemical equilibrium model and coupled fluid flow-precipitation kinetics model is described below. As noted already, Kuroko deposits are characterized by the following zonal arrangement in ascending stratigraphic order: siliceous ore (quartz, chalcopyrite, pyrite), yellow ore (chalcopyrite, pyrite), black ore (sphalerite, galena, barite), barite ore (barite and quartz) and ferruginous chert ore (microcrystalline quartz, hematite). Quartz is abundant but barite is poor in the siliceous ore, though barite veinlets occur in this zone. Barite is common in the black ore and abundant in barite ore. Barite is also found in ferruginous chert ore (Kalogeropoulos and Scott, 1983). Quartz is poor in massive sulfide ore horizons (yellow and black ores). The distribution of quartz and barite in Kuroko deposits suggests that barite and quartz precipitate by different mechanisms. Figure 1.50 shows the relationship between barite and quartz contents of the Kuroko ore samples. This suggests that quartz and barite formed separately in different parts (Ohmoto et al., 1983). Amorphous silica and barite precipitate simultaneously from white smoker in midoceanic ridge hydrothermal system (Edmond et al., 1979). It is inferred that amorphous silica precipitates in the chimney at a later stage than sulfides and sulfates (anhydrite and barite) which constitute chimneys from which black smoker is emerging. It is thought that the precipitation of amorphous silica is caused by conductive cooling from the hydrothermat solution which flows laterally in the chimney (Herzig et al., 1988). Barite is abundant in back-arc basin hydrothermal system such as Okinawa, Manus and Mariana (Shikazono and Kusakabe, 1999). In these chimneys, coprecipitation of barite and amorphous silica is taking place from the solution characterized by lower temperatures and lower flow rate than the black smoker. Solubilities of quartz and amorphous silica in aqueous solutions increase with increasing of temperature (Holland and Malinin, 1979). Solubility of barite depends on salinity and temperature (Blount, 1977). The solubility of barite in hydrothermal solution having more than 1 molal NaC1 concentration increases with increasing temperature, while a solubility maximum exists in the solution with NaC1 concentration less than ca. 0.2 molal (Blount, 1977).
Chapter 1
68
==
./
<
om
0
1O0
200
300
400
500
600
Precipitated Amount of Quartz (mg/kg • H20) Figure 1.51. Relationship between precipitated amount of quartz and that of barite.
Salinity of ore fluids responsible for Kuroko deposits is in a range of 0.5-1 tool/1 which is estimated based on the fluid inclusion studies (e.g., Marutani and Takenouchi, 1978). Thus, it seems likely that the precipitation of barite takes place by decreasing of temperature and increasing of concentration of sulfate ion caused by the mixing of hydrothermal solution having 0.5-1 molal NaC1 with cold ambient seawater. Quartz precipitates also due to the mixing of hydrothermal solution with ambient cold seawater if equilibrium between solution and quartz is attained. The thermochemical equilibrium calculations on the amounts of minerals precipitated due to the mixing of ascending hydrothermal solution with cold seawater have also shown that the amounts of quartz precipitated correlate to that of barite precipitated (e.g., Ohmoto et al., 1983; Janecky and Seyfried, 1984; Bowers et al., 1985). Previous studies on the precipitations of barite and quartz calculated the saturation index with respect to barite and quartz during the mixing of hydrothermal solution and seawater. However, the amounts of barite and quartz precipitated were not obtained. Therefore, the precipitated amounts due to the mixing was calculated. The relationship between the amounts of barite and quartz precipitated is shown in Fig. 1.51, which indicates that these amounts are positively correlated. However, as ah'eady noted, the barite content in Kuroko ore inversely correlates to the quartz content and the occurrences of barite and quartz in the submarine hydrothermal ore deposits are different. The discrepancy between the results of thermochemical equilibrium calculations based on the mixing model and the mode of occurrences of barite and quartz in the submarine hydrothermal ore deposits clearly indicate that barite and quartz precipitated from supersaturated solutions under non-equilibrium conditions. Thus, it is considered that the flow rate and precipitation kinetics affect the precipitations of barite and quartz. It was attempted to derive the relationships in the precipitated amounts of barite and quartz, flow rate and precipitation rate using the coupled fluid flow-precipitation
Miocene-Pliocene Hydrothermal Ore Deposits
69
nit J
I
Mixing
Precipitation
/
q C, Figure 1.52. Uniformly mixed kinetics model, q: volume flow rate (m3/s), Ci: initial concentration (molal), l: height (m), r: radius (m), C: concentration (molal), V: volume of the system (m3). model. T w o g e o l o g i c sites o f the precipitations o f barite and quartz are taken into account: one is the site a b o v e the seafloor and the other one is that b e l o w the seafloor. The calculations w e r e m a d e based on the u n i f o r m l y m i x e d precipitation kinetics m o d e l (Fig. 1.52) for one c o m p o n e n t and one d i m e n s i o n a l system r e p r e s e n t e d by Eq. (1-20).
dC / dt = - k ( a / m ) ( c - Co) + (q / V) ( Ci -- C)
(1-20)
w h e r e k = precipitation rate constant, A = surface area on w h i c h barite and quartz precipitate, M = mass o f h y d r o t h e r m a l solution, C = c o n c e n t r a t i o n o f Ba 2+ and H4SiO4 o f the solution in the system, Ci: the c o n c e n t r a t i o n o f input solution, q = v o l u m e flow rate, and V = v o l u m e o f the system. C a l c u l a t e d results are shown in Figs. 1.53 and 1.54. A s s u m e d values o f parameters (l, r, A / M ) for the calculations are g i v e n in Table 1.8. A c o m p a r i s o n o f the calculated results (Figs. 1.53 and 1.54) with the m o d e o f o c c u r r e n c e s o f quartz and barite in the s u b m a r i n e h y d r o t h e r m a l ore deposits indicates
TABLE 1.8 Parameter values used for the computations l (m)
r (m)
A/M (m2/kg)
q/V (l/s)
1
1
1 0.1
0.01 0.1 0.1
0.2 0.02 0.02
v
2 3
v 10 v
v: velocity (m/s); A: surface area (m2); M: mass of fluid (kg); q: volume flow rate (m3/s); V: volume of system (m3).
Chapter i
70
0.16
I
0.12
----
l=lm,r=0.01m l=lm,r=0.1m I I--0.1m,r=-0.1m CBai
O
<
~.~ 0.08
X \ \
0.04
0
\
~"
-,,
~
. . . i . . . .T . . . " r '.- . .---..-t. . . . i,. . . .¢--..
i
2
3
4
5
6
7
8
9
10
Velocity(m/s) Figure 1.53. Relationship between precipitated amount of barite and velocity of fluid at 200°C. v: velocity (m/s), CBai: initial concentration of Ba. 14
12
10
l=Im,r=O.Olm I l=Im,r=O.Im I l=O.Im,r=O.Im Cs o2i
- - - -
81',` "~
4
21| 0 1
-8
\
\",,~
\\\ \ i" - ~
-7
-'-'~...,-,---
-6
- ".~. ~ - -
-5
--~
. . . . .
-4
L------.J
-3
. . . .
-2
-1
Log 1~ Figure 1.54. Relationship between precipitated amount of quartz and velocity of fluid at 200°C. Csio2i: initial concentration of SiO2; v: velocity (m/s).
Miocene-Pliocene Hydrothermal Ore Deposits
71
that the precipitations of barite and quartz take place in different sites in the submarine hydrothermal system. Quartz tends to precipitate under the conditions of high temperature, high A/M, and slow fluid flow rate, while barite under the conditions of low temperature, low A/M, and high fluid flow rate. These results seem to be in agreement with the mode of occurrence of barite and quartz. Barite in the black ore is thought to have formed from the fluids with high flow rate (1-10 m/s) on the seafloor. Barite is found in the chimney from which black smoker is venting. On the other hand, quartz formed in the siliceous ore in high A/M, replacing the dacite below the seafloor. Ferruginous chert in which abundant silica occurs formed below the seafloor by the mixing of ferruginous sediments and hydrothermal components (Kalogeropourous and Scott, 1983). Barite-silica chimney found in back-arc basin formed in the conditions similar to that of ferruginous chert and barite bed in the Kuroko deposits; temperature is relatively low (ca. 150-100°C), and flow rate of fluids may be slow. The above mentioned study on barite and quartz precipitations in Kuroko deposits is summarized as follows. (1) Quartz occurs abundantly in feeder ore (stockwork siliceous ore) in Kuroko deposits. (2) Quartz coexisting with barite also occurs in the ferruginous and barite ores in Kuroko deposits. (3) Barite is abundant in the massive strata-bound ore bodies (black and barite ores) in Kuroko deposits and occurs in the ferruginous chert ore in Kuroko deposits, and chimneys in active deposits at back-arc basins. The coupled fluid flow-precipitation kinetics model calculations indicate the following results: (1) Quartz or amorphous silica tends to precipitate from the solution having relatively high temperature and low flow rate and under high A/M condition. (2) Barite tends to precipitate from the solution with relatively high flow rate and low temperature and under low AIM. (3) These predictions are generally in agreement with the observations; homogenization temperatures of fluid inclusions in quartz from siliceous ore zone and in barite from black ore zone in the Kuroko deposits is relatively high, ranging from 350 to 250°C, and tow, ranging from 250 to 150°C, respectively. (4) The results of calculations are in agreement with the occurrences of barite and silica and chemical features of discharging fluids in the submarine hydrothermal ore deposits; namely, quartz is inferred to precipitate in subseafloor environment and barite in seabottom environment. A/M for the stockwork siliceous ore zone in Kuroko deposits is low, while that for black ore in the Kuroko deposits and chimney may be high. The above-mentioned consideration indicates that important factors controlling the precipitations of barite and silica are surface area/water mass ratio (A/M), temperature, precipitation rate constant (k) and flow rate (v), and the coupled fluid flow-precipitation models are applicable to understanding the distributions of minerals in submarine hydrothermal ore deposits.
Chapter I
72
Barite precipitation highly depends on SO ] . and Ba 2+ concentrations in the fluids. That means that the mixing ratio of hydrothermal solution and seawater is also an important factor for the precipitation of barite, together with the factors mentioned above.
Precipitation mechanism of barite. The precipitation mechanism of barite from aqueous solutions at temperatures below 100°C have been extensively studied by many workers (e.g., Elving and Leineweber, 1950; Wagner and Wuellner, 1952; Nielsen, 1955, 1957, 1958, 1959a,b, 1961; Nielsen and Tort, 1984; Collins and Leineweber, 1956; Takiyama, 1959a,b; Fabrikanos and Lieser, 1962; Lieser and Wertenbach, 1962; Nancollas and Liu, 1975; Klein and Frontal, 1964; Blount, 1974; Liu et al., 1976; Rizkalla, 1983). However, barites in back-arc basin deposits precipitated in the temperatures more than 100°C, such as 150-300°C in Kuroko deposits (e.g., Pisutha-Arnond and Ohmoto, 1983). Thus, Shikazono (1994) carried out the experiments of barite precipitation at elevated temperature (150°C). It was found from his experiments that the morphology of barite varied with the barium chloride concentration. Dendritic crystals with rod-like, spindlelike, and star-like or cross-like habits, and irregularly shaped crystals with rough surface formed from solutions of high barium chloride concentrations (BaCl2, 0.08-0.8 molal). Feather-like dendritic crystal did not form. The results obtained by Shikazono (1994) are generally in agreement with those obtained by previous investigations in their barite precipitation experiments at less than 100°C. For example, Liu et al. (1976) found that dendritic crystals of barite formed at 25°C from aqueous solution containing BaSO4 higher than 2.0 × 10 - 3 molal, which is in agreement with the result by Lieser and Wertenbach (1962). They synthesized feather-like barite from aqueous solutions containing 2.0 x 10.2 molal BaSO4. Suito and Takiyama (1954) and Takiyama (1959a) reported that spindle-shaped dendritic barite crystals having ragged edges formed from aqueous solutions having 1.0 x 10 2 to 2.0 x 10-2 molal BaSO4. Sasaki and Minato (1983) reported that barites precipitated at 60°C from an aqueous solution having the total cation (Ba(NO3)2 + Pb(NO3)2) concentration greater than 0.004 molal were very fine-grained, and complex dendritic particles crystals, while rectangular (Ba,Pb)SO4 crystals formed from a solution of lower concentration. From the values of BaCl2 concentrations and the dissociation constant of BaC1+ (Helgeson and Kirkham, 1976), the concentrations of Ba 2+ in the experimental solutions are estimated to be 0.034-0.047 molal. Assuming the above values of SO42 and Ba 2+ concentrations, the boundary condition for the formation of dendritic barite and wellformed barite is estimated to be - 6 . 2 to 6.0 for the log(mBa2+)i(mso2)i value at 150°C (Fig. 1.55). The log(mBa2+)i(mso2-)i values for the solutions of the previous investigators (T = 25-100 C) and observed morphologies of barite m their experiments) are compared as shown in Fig. 1.55. It shows that the morphology of barite crystals changes with an increase in the concentration product, (mBa2+)i(mso2)i, from well-formed (rectangular, rhombohedra) (R) through rod-hke, star-hke, and spmdle-hke dendrmc (D) to feather-hke dendritic crystals (DF). The concentration products obtained in the experiment by Shikazono (1994) can be compared with the solubility product for large well-formed polyhedral barite crystals o
4
.
•
.
.
.
.
,4
.
.
.
.
.
.
Miocene-Pliocene Hydrothermal Ore Deposits -2
DF
-3
DF
-4
DF
73
D D D D
0-~
D D
m 0"~-8
0 __1
-9 -10 [
I
t
~o
,oo
,5o
i
200 V (°(3)
Figure 1.55. The relationships between the concentration product, (Ba2+)i(SO42-)i,at the initiation of barite precipitation, and morphologies of barite crystals (Shikazono, 1994). The dashed line represents the boundary between dendritic barite crystals and well-formed rhombohedral, rectangular, and polyhedral barite crystals. The 150°C data are from Shikazono (1994); the others from other investigations. D: dendritic (spindle-like, rodlike, star-like, cross-like) barite; DF: feather-like dendritic barite; W: well-formed rectangular, rhombohedral, and polyhedral barite. (i): The boundary between the diffusion-controlled mechanism (Di) and the surface reaction mechanism (S) for barite precipitation at 25°C estimated by Nielsen (1958); ®: The solubility product for barite in 1 molal NaC1 solution at I50°C based on data by Helgeson (1969) and Blount (1977). A-B: The solubility product for barite in 1 molaI NaCI solution from 25 to 150°C based on data by Helgeson (1969).
adopted by Helgeson (1969) and experimentally obtained by Blount (1977). The boundary determined by Shikazono (1994)'s study (log(mBa2+)i(mso2)i = - 6 . 1 ± 0 . 1 at 150°C) 4. lies above the concentration products in 1 molal NaC1 solunon by Helgeson (1969) and Blount (1977) which are plotted as two (the concentration product based on the Blount (1977) experimental results is - 7 . 3 in log unit) for 150°C) and the curve A - B from 25 to 150°C (Fig. 1.55). The above comparison indicates that the solubility of dendritic barite crystals is higher than that of well-formed polyhedral large barite crystals used by Blount (1977). The saturation index for the boundary between dendritic barite and well-formed barite can be estimated to be 10 - 6 ' I / 1 0 .7.3 which is equal to ca. 20. It is usually believed that the growth o f dendritic crystals is controlled by a bulk diffusion-controlled process which is defined as a process controlled by a transportation of solute species by diffusion from the bulk of aqueous solution to the growing crystals (e.g., Strickland-Constable, 1968; Liu et al., 1976). The appearances of feather- and star-like dendritic shapes indicate that the concentrations of pertinent species (e.g., Ba 2+, SO 2 - ) in the solution are highest at the corners of crystals. The rectangular (orthorhombic) crystal forms are generated where the concentrations o f solute species are approximately the same for all surfaces but it cannot be homogeneous when the consumption rate o f solute is faster than the supply rate by diffusion (Nielsen, 1958).
Chapter 1
74
The dependency of rate of precipitation of barite from aqueous solution on time at room temperature studied by Nielsen (1958) suggests that the precipitation of barite from solutions of high levels of supersaturation (i.e., more than 30 as saturation index (S.I.) which is defined as the ratio, (mBa2+)(mso2-)/Ksp, where Ksp is the solubility product m molahty for equlhbrlum) is controlled 15y a bulk diffusion mechanism, while the surface reaction mechanism (polynuclear growth) dominates at low S.I. (i.e., less than 30). The boundary between the bulk diffusion mechanism and the surface reaction mechanism (polynuclear growth mechanism), suggested by Nielsen (1958) at 25°C, is shown in Fig. 1.55 ((~ in the figure). His estimated boundary is consistent with that between dendritic and well-formed barite which was determined by scanning microscopic measurements by other workers (e.g., Fisher and Rhinehammer, 1953; Okada and Magari, 1955; Liu et al., 1976). Nielsen and Toft (1984) summarized the relationship between precipitation mechanism of sparingly solute electrolytes (e.g., BaSO4, AgC1, CaF2) and the degree of supersaturation, distinguishing the surface reaction-controlled and diffusion-controlled regions by plotting growth rates of barite on a PBa-PSO4 diagram (where PBa log EBa; EBa is total dissolved barium concentration; Pso4 log ESO4, a total dissolved sulfate concentration). Nancollas and his group (e.g., Nancollas and Liu, 1975) have suggested second order kinetics for barite precipitation at 25°C with a surface-controlled mechanism even at an S.I. as high as 56. These studies have demonstrated that the surface reaction mechanism dominates to the S.I. of ca. 10-100 and that above this S.I. the bulk diffusion mechanism controls the precipitation of barite. This boundary between two different precipitation mechanisms for barite determined by the experiments at 150°C (Shikazono, 1994) roughly coincides with that between dendritic crystals and well-formed crystals which has been experimentally determined at temperatures lower than 100°C. Several morphologies of dendritic barite, such as feather-like, rod-like, spindlelike, star-like, and cross-like crystals have been recognized when the (mBa2+)i(ms02-)i values were considerably higher than the equilibrium values, although no detailed studies have been made on the relationship between the various morphologies of dendritic barites and the degree of supersaturation. Barite is abundant and widespread in Kuroko deposits. However, it is concentrated especially in the upper horizons (black ore and barite ore), increasing upwards within the black ore. Such a trend is also observed frequently in the tetsusekiei, but rarely in the yellow ore. Barite occurs also as vein-fillings in the stockwork siliceous ore. There are four main types of barite occurrence. Type A is well-formed coarsegrained barite in the black and yellow ores, associated with sulfide minerals (sphalerite, galena, chalcopyrite, pyrite). Fine-grained barite (type B) is also associated with sulfides in the black and yellow ores. Fine-grained barite tends to occur in the upper parts of black ore and also in yellow ore. According to Eldridge et al. (1983), fine-grained sulfides and barite were primitive, rapidly precipitated due to the mixing of ascending hot solution with cold seawater at the seafloor (Sato, 1973; Eldridge et al., 1983). They converted to the coarser-grained barite through dissolution and recrystallization within the ore piles. Fine-grained barite tends to occur in black ore, while coarse-grained barite is found in the .
.
.
.
.
.
4
=
=
- -
- -
Miocene-Pliocene Hydrothermal Ore Deposits
75
lower horizons of black ore and also in yellow ore. Type C is fine-grained barite in the massive barite ore horizon. Under an optical microscope, type C appears as an aggregate of radial crystals, intergrowing with fine-grained quartz. Type C is more elongated in shape compared to types A and B. Most barites (types A, B, and C) are rectangular, lath-shaped, platy, and tabular. Rare type D barite crystals are rectangular and polyhedral. Polyhedral, rhombohedral, or coarse-grained barite crystals (type D) are rarely observed in vugs of black ore or in the inner parts of the chimney recovered from the Hanaoka mine (Shimazaki and Horikoshi, 1990; Shikazono, 1992). The various barite types were examined by scanning electron microscopy (SEM). No dendritic (feather-shaped, spindle-like, rod-like, star-like, or cross-like) barite was found. Coarse-grained barite (type A) and fine-grained barite (type B) have generally rectangular shape but sometimes irregular shapes with rough surfaces. The surface of type B is rougher than that of type than that of type A. Type A and B barites are comprised of very fine-grained, disk-like barite microcrystals. The grain size of an individual microcrystal ranges from ca. 0.1 to 5 Ixm. Massive barite crystals (type C) are also composed of very fine grain-sized (several txm) microcrystals and have rough surfaces. Very fine barite particles are found on outer rims of the Hanaoka Kuroko chimney, while polyhedral well-formed barite is in the inner side of the chimney (type D). Type D barite is rarely observed in black ore. These scanning electron microscopic observations suggest that barite precipitation was controlled by a surface reaction mechanism (probably surface nucleation, but not spiral growth mechanism) rather than by a bulk diffusion mechanism. There are two interpretations for the formation of fine disk-like barite particles which constitute types A, B, and C barites. One is homogeneous nucleation caused by rapid mixing of hydrothermal solution with cold seawater and the coagulation of resulting fine particles. The other is heterogeneous nucleation and growth of barite on the surfaces of pre-existing barite crystal. Calculations of settling velocities for very fine particles in hydrothermal plumes issuing from the ocean floor suggest that very fine (less than 10 g m in diameter) particles cannot settle onto the nearby ocean floor (Shikazono, 1992). However, barite microparticles with grain sizes less than 10 ~ m are common in Kuroko deposits. Therefore, heterogeneous nucleation is likely rather than homogeneous nucleation. Sulfate-sulfide chimneys have been recently discovered from the Mariana back-arc basin (e.g., Kusakabe et al., 1990). They are composed of barite, sphalerite, galena, chalcopyrite, and silica. Barite is most abundant in the chimney. Scanning electron micrographs of barites in a Kuroko chimney show that they are well-formed polyhedra. Morphologically similar barite is also found in the inner side of a chimney from the Mariana back-arc basin. Such barites appear to have formed by recrystallization of fine barites which occur in the outer side of the chimneys. The various morphological features of barites from the Kuroko and Mariana deposits, when combined with the experimental studies on barite precipitation, suggest that the surface reaction mechanism was dominant for the formation of these barites. This implies that the concentration product, (mBa2+)(mso2-), at the initiation of barite precipitation was probably less than ca. 100 times that for equilibrium. This estimate can be evaluated based on the strontium and sulfur isotopic studies
-..-1
TABLE 1.9 Chemical composition of hydrothermal solution experimentally interacted with rocks
T (°C) P (bar} W/R Day's pH Na K Ca Mo Fe Mn Si AI CO2 SO4 H2S Cu Ni Zn Pb Au Hg As Cd Sb Ba Sr
1 (ppm)
2
3
4
5
6
7
8
9
10
11
12
13
14
350
350
455 764 349 3.7 14.9 288 0.04 1475 250 <0.1 0.02 0.3 l 0.17 0.01 0.017 0.02 0.7 0.02 0.04 1.2
1561 1330 2 5.6 8.8 1154 0.05
202 1,5 10.5 8.9 184(I 4.1
4(I(I 1000 5 14 3.7 10750 2300 313 9.0 62.5 54 1275
500 1000 5 14 3.0 10250 460(1 370 5 490 76 1200
500 1000 5 ~4 1.0 11250 950 310 250 7.1 7.1 1500
500 1000 5 14 3.5 8250 1860 1530 14.5 1100 120 1300
500 1000 5 14 3.5 10300 425 1100 7.6 385 70 1100
300 700 3 236 5.35 11084 726 1856 5.6 2.0 2.7 416 <0.26 159 1.39 6.60 <0.05 <0.07 <0.15
400 700 3 ~00 3.85 l 1156 762 1565 10.5 117 53 1684 <(}.26 573 5.65 7.54 <(}.05 <0.07
262 500 50 100 4.65 10300 370 1420 351 45 110 1350 0.21
-
10 16 4.8
300 10(,10 5 t4 3.0 750(I 1150 96 345 25.5 4.6 650
300
10 124 5.2
300 1000 09 16 4.2
2.0 5.5
2.1 6.9
10 40 0.11 0.7(I 1.5 (1.05 0.052 0.06 0.6 0,09 0.08 2.1
1.5
4.1
1 2: graywacke; 3-7: rhyolite: 8: andesite; 9-13: basalt; 14: seawater.
1.1
92 5.6 10963 489 818 0.l 3.5 0,05 580 0.37 118 20 10 0,01 0.01 0.1
0,135
0.50
0.38
<0.15
7.89 10760 399 412 1294 2 x 10 3 0.2 x 10 -3 2900 x 10 3 2 x 10 - 3 145 2717 0.3 x 10 3 0.6 x 10 3 3 x 10 -3 0.03 × 10 -3 0,005 × 10-3 0.1 x 10 3 3 x 10 _3 0.1 x 10 -3 0.33 x 10 -3 20 x 10- 1 7.9 ~"
Miocene-Pliocene Hydrothermal Ore Deposits
77
of barite and strontium content of anhydrite from Kuroko and Mariana deposits by Shikazono et al. (1983), Kusakabe and Chiba (1983), and Kusakabe et al. (1990). These authors indicated that these sulfate minerals precipitated due to the mixing of the hydrothermal solution with cold seawater with a contribution of the end member hydrothermal fluid to seawater being less than 20%. Ohmoto et al. (1983) calculated S.I. of barite to be less than ca. 101'5 during the mixing and cooling of fluids. The estimated value is in agreement with that of Shikazono (1994).
1.3.4.2. Origin of ore fluids Origin of ore fluids is constrained by (1) chemical compositions of ore fluids estimated by thermochemical calculations (section 1.3.2) and by fluid inclusion analyses, (2) isotopic compositions of ore fluids estimated by the analyses of minerals and fluid inclusions (section 1.3.3), (3) seawater-rock interaction experiments, (4) computer calculations on the seawater-rock interaction, and (5) comparison of chemical features of Kuroko ore fluids with those of present-day hydrothermal solutions venting from seafloor (section 2.3). During the last two decades, many experimental studies on the seawater-rock interaction at elevated temperatures (100~00°C) have been conducted. Particularly, detailed seawater-basalt interaction experiments have been done. Several experimental studies on seawater-rhyolite interaction and seawater-sedimentary rock interaction are also available (Bischoff et al., 1981). Examples of chemical compositions of modified seawater experimentally interacted with various kinds of rocks are shown in Table 1.9. Several factors such as C1- concentration, water/rock ratio and temperature are important in controlling the chemical composition of the hydrothermal solution interacted with the rocks. For example, water/rock ratio affects the alteration mineralogy (Mottl and Holland, 1978; Seyfried and Mottl, 1982; Shikazono, 1984). For example, at low water/rock ratio, epidote is stable, while chlorite at high water/rock ratio (Shikazono, 1984; Shikazono and Kawahata, 1987). From rock-water interaction experiments (Table 1.9) and analytical data on fluid inclusions we can derive the relationship between the concentrations of alkali, alkali earth and base-metal elements and concentration of C1- ion in the hydrothermal solutions experimentally interacted with rocks, in the natural hydrothermal solutions and in fluid inclusions (Figs. 1.56-1.58). It is seen in Figs. 1.56-1.58 that the concentrations of elements increase with increasing of C1- concentration. Especially base-metal (Zn, Fe, Mn, Pb) and Ba concentrations increase rapidly with increasing of C1- concentration. This relationship strongly suggests that C1- concentration is a very important factor controlling the chemical compositions of ore solution. Detailed discussion on the relationship between C1- concentration and concentrations of alkali and alkali earth elements were carried out by Shikazono (1978a) (see section 2.1). In Fig. 1.59 the relationship between temperature and concentration of elements (Zn, Ba) at constant C1- concentration which is equal to that of seawater obtained by the experimental studies and analytical data on natural hydrothermal solution (geothermal water) are shown. It is seen that the concentrations of base-metal elements (Zn, Fe, Mn, Cu, Pb) and Ba increase with increasing of temperature. Concentrations of these
Chapter 1
78
Log(Zn) 3 OOX
2 1 x
0
x
x •
x x
x
-;
x x
-3
x
-Z
:;
o
:3
4
,5 Log(CI)
Figure 1.56. Relationship between the zinc and CI- concentration in geothermal waters and hydrothermal solution experimentally interacted with rocks (Shikazono, 1988c).
Loq(Fe)
,"
3
:
," o
,'* ;, , II
G
,~I ' /s " X
1
•
,
st 0
~
o0
,,~f
," X
-2
o0
:2[ -3
0
1
X
~. sesX
sS
-1
,o . ,
!
x, " "
• S
""
•
," e °
oO S.W
:;
:3
4
5
Log(Cl)
Figure 1.57. Relationship between the iron and C] concentration in geothermal waters and hydrothermaI solution experimentally interacted with rocks (Shikazono, 1988c).
Miocene-Pliocene Hydrothermal Ore Deposits
79
Log (Ba) CI = S.W. 4 3 AR. 2
% %
gR.
1
~X
0
X
0
%
-1
% %
% %
-2
% %
% % %
-3
SW
-4
800
1000
600
400
200
T(°C)
0
Figure 1.58. Relationship between Ba2+ concentration and temperature of geothermal waters and hydrothermal solution experimentally interacted with rocks (AR: acidic rocks, BR: basic rock) (Shikazono, 1988c).
Loq(Zn) CI = S.W. 5 4 3 I
2
6
%
1
%
~X
0
% ,%
-1
;-. X"
-2
% %
x
SW
-31000
I
I
800
I
I
600
i
1
400
I
I
200
,
I
o
T(°C)
Figure 1.59. Relationship between zinc concentration and temperature of geothermal waters and hydrothermal solution experimentally interacted with rocks (Shikazono, 1988c).
elements at hydrothermal conditions (200-400°C) are much higher than those of seawater. Therefore, it is evident that the effect of temperature on the concentrations of ore-forming elements (base metals, Ba) is strong. For example, the concentrations of base metal elements in modified seawater interacted with acidic volcanic rocks at 300°C are Fe, 100
Chapter 1
80
TABLE 1.10 Isotopic compositions of Kuroko ore solution (K.O.), seawater (S.W.) and magmatic water (M.W.) (Shikazono, 1978) K.O.
S.W.
M.W.
3348E,o,S ~348I],r,S ~)34Stota1
+21 tO +24%o +2 to +5%o >+2 tO +5%o
+20%o +20%0
0 to +1%,)
~180 ~D 87Sr/86Sr
0 to +5%0 to 30 to -3%0 0.7062-0.7087
0%o 0%o 0.709
+7 to +9%° --48 to --85% 0.704
334Sz,o,S:Sulfur isotopic composition of oxidized sulfur species. ~34S~,r,S: Sulfur isotopic composition of reduced sulfur species. ~348total:Total sulfur isotopic composition of aqueous sulfur species. ppm; Ba, 10 ppm; Zn, 1 ppm; and Cu, 0.5 ppm (Table 1.9). These concentrations are roughly in agreement with those estimated by solubility calculations on Kuroko ore fluids (Tables 1.5 and 1.6). Isotopic compositions of minerals and fluid inclusions can be used to estimate those of Kuroko ore fluids. Estimated isotopic compositions of Kuroko ore fluids are given in Table 1.10. All these data indicate that the isotopic compositions lie between seawater value and igneous value. For instance, 87Sr/86Sr of ore fluids responsible for barite and anhydrite precipitations is 0.7069~).7087, and 0.7082-0.7087, respectively which are between present-day seawater value (0.7091) and igneous value (0.704~3.705). From these data, Shikazono et al. (1983), Farrell and Holland (1983) and Kusakabe and Chiba (1983) thought that barite and anhydrite precipitated by the mixing of hydrothermal solution with low 87Sff86Sr and seawater with high 87Sr/86Sr. ~34S values of sulfides are +2%o to +7%o, which lie between igneous value (0%o (typical value)) and seawater value (+20%o). This suggests that both seawater sulfur and sulfide sulfur of the igneous rocks were incorporated into the Kuroko ore fluids of modified seawater origin. Probably, the sulfide sulfur of Kuroko ore fluids were derived from partial reduction of seawater sulfur and dissolution of pyrite and anhydrite in the country rocks. It is noteworthy that 334S values of sulfides from small B sub-type are smaller than large C sub-type and Y sub-type. This difference could be explained in terms of the extent of seawater sulfate reduction to H2S. The above argument on the calculation of chemical composition of ore fluids, seawater-rock interaction experiments, and isotopic compositions of ore fluids clearly demonstrates that Kuroko ore fluids were generated by seawater-rock interaction at elevated temperatures. The chemistry of present-day hydrothermal solution venting from back-arc basins and midoceanic ridges (sections 2.3 and 2.4) also support this view. There is another opinion on the origin of Kuroko ore fluids. Sawkins (1982) thought that intrusive felsic magmas were the source of the metals and heat in Kuroko hydrothermal systems. He stressed the contributions of magmatic fluid and seawater in
Miocene-Pliocene Hydrothermal Ore Deposits
81
ore fluids from which Kuroko deposits formed based on 3D, 3180 (Hattori and Sakai, 1979; Hattori and Muhlenbachs, 1980) and salinity data on inclusion fluids. His calculated mixing ratio of magmatic fluid to seawater was 1/9. The problem on his magmatic model is that metal content of magmatic fluid mixed with cooler hydrothermal solution of seawater origin is high and metal sulfides tend to precipitate even under the small mixing ratio of magmatic fluid to hydrothermal solution probably at very high temperatures. Another problem of magmatic hypothesis which has been critically thought by Ohmoto et al. (1983) is sulfur behaviour during this type of mixing. High salinity of Kuroko ore fluids does not solely mean magmatic contribution. Instead salinity variation can be reasonably explained by subcritical boiling of fluids of seawater origin. ~180 and 3D data (Fig. 1.40) do not clearly show the mixing of seawater and magmatic water which was favoured by Ishihara and Sasaki (1978). Ohmoto et al. (1983) pointed out that the first problem of the magmatic hypothesis is the large difference in the S concentrations between Kuroko ore fluids and magmatic fluids; fluids derived from acidic magmas at temperatures >750°C are more likely to contain ~ 1 0 -1 to > 1 M ~2S (Burham and Ohmoto, 1980), whereas Kuroko ore fluids appear to have contained 10-3-10 -2 M ZS at 250 4-50°C. The second and most serious problem pointed out by Ohmoto et al. (1983) is that magmatic fluids with +5%e sulfur which was estimated based on 334S values of granitic rocks in Green tuff region by Ishihara and Sasaki (1978) are unlikely to form sulfides of ~+5%o at temperature <300°C under the chemical conditions of the Kuroko ore fluids. Even if we accept that ~D and 3180 values lie between the two end member ~ values, this feature can be also explained by low ~D and high 3180 solution being generated by low seawater/rock ratio condition. Lead isotopic data support this interpretation; namely, these data clearly indicate leaching of lead from the rocks (Fehn et al., 1983). Horikoshi and Shikazono (1978) indicated that 3D of ore fluids for B sub-type which is located at centre of Hokuroku basin is higher, suggesting large contribution of seawater, while 3D of ore fluids of Y sub-type located at the margin of Hokuroku basin is lower, suggeting meteoric water contribution. Ohmoto et al. (1983) thought that the involvement of meteoric water in Kuroko ore fluids is unlikely because they estimated the entire Hokuroku district was under more than 2,500 m of seawater depth (probably ca. 3,500 m and the district was located more than 50 km away from the nearest island based on the studies of foraminifera assemblage in mudstone overlying Kuroko deposits (Guber and Ohmoto, 1978; Guber and Merrill, 1983). However, the seawater depth is inferred to be shallower probably 1,000-2,000 m from the following reasons. One of the reasons is that the seawater depth of hydrothermal ore depositions recently discovered at back-arc basins (Okinawa Trough, Izu-Bonin, Mariana-Trough, Marius Basin, North Fiji Basin, Lau Basin) is mostly 1,000-2,000 m. Oki and Hayasaka (1978) reported an occurrence of recent arenaceous foraminifera in a shallow submarine caldera, which usually lies at a depth of more than 4,000 m on the bottom of the open ocean. The seawater in the caldera has a lower pH than normal seawater due to the inputs of volcanic gas and hydrothermal solution with low pH. Lower pH raises the carbonate compensation depth (CCD) there. Therefore, the estimate of
Chapter 1
82
E v
o)
7-]:
20
19
18
3"~ ~ 17
I ~.~
I 16
15
time
14
(Ma)
Figure 1.60. Variation of subsidence rate for syn-rift basins in the Uetsu district, northeast Honshu (Yamaji, I990). The line of boxes shows the spatially averaged subsidence rate. The rate after 15 Ma is not clear because of uncertainty in paleobathymetry. However, the rate probably decreased to the order of 10-100 m/m.y. If the rate had been of the order of 1 km/m.y, after 15 Ma, the water depth of the inner arc region at 14 Ma would have been much deeper than modem, young, back-arc basins.
seawater depth based on paleobathymetry using CCD as a reference point in the Miocene open ocean by Guber and Ohmoto (1978) might be invalid (Kitazato, 1979; Matoba, 1983; Yam@, 1990). Kitazato's (1979) estimate of 1,000-2,000 m as the seawater depth of Kuroko formation seems more likely. Akimoto and Hasegawa (1989) also estimated to be 1,000-2,500 m. Their estimated seawater depth is similar to that of present-day back-arc basins from which hydrothermal ventings occur (Okinawa Trough, Mariana). Yamaji (1990) studied paleobathymetry of the Northeast Japan Arc at 17-14 Ma (Fig. 1.60) and showed that the arc subsided rapidly from 16 to 15 Ma (Fig. 1.61). According to his paleobathymetry, it is likely that shallow sea (sublittoral ~ 1 5 0 m and upper bathyal 150-500 m) existed near the site of Kuroko formation at 15-16 Ma. Considering the discussion above, it cannot be ruled out that meteoric water was involved in the Kuroko ore fluids. In fact, ~D and 3180 values of fluid inclusions from Iwami Kuroko deposit, west Honshu (San-in district) are considerably lower than those of seawater, indicating an involvement of meteoric water. However, 3D and 3180 values of Nurukawa Kuroko ore fluids (Y sub-type) are plotted between seawater value and igneous (or magmatic) value (Ishiyama et al., 2001), suggesting an involvement of igneous (or magmatic) fluids and no contribution of meteoric water. Here, igneous fluids mean the fluids controlled by igneous rocks, which were generated under a low water/rock ratio.
Miocene-Pliocene Hydrothermal Ore Deposits
®
83
Q •
S
SS
S
/~'Ooo s~ ,~
N = NON-MARINE SUBLITTORAL (0-150 m) o = UPPER BATHYAL (150-500 m) • = M I D D L E T O L O W E R BATHYAL (500-2500 m) S =
,
100 km
Figure 1.61. Paleobathymetry of the NE Japan Are at 17-14 Ma (circled numbers) (Yamaji, 1990). The arc subsided rapidly from 16 to 15 Ma. The entire inner arc region became at middle bathyal depths until 15 Ma, except that Yuri hill remained at shallow marine depths. The hill was submergeduntil 14 Ma (Matoba, 1981). Stippled area, Aosawa basalt (Tsuchiya, 1988). Hatched area, Yuri hill. Note that submarine volcanoes are neglected in this figure because volcaniclasticrocks usually contain few fossils.
The most important conclusion derived from the isotopic studies mentioned above is that isotopic characteristics of Kuroko ore fluids were caused dominantly by seawater-volcanic rock interaction at elevated temperature and by the mixing of seawater with small portions of igneous water or the hydrothermal solution whose chemical and isotopic compositions are controlled by water-rock interaction under the rock-dominated condition and also small proportion of mixing of meteoric water. However, it cannot be decided at present which processes (degree of seawaterrock interaction or mixing ratio of seawater, igneous water and meteoric water) are important for the generation of Kuroko ore fluids solely from the isotopic studies. But experimental and theoretical considerations on seawater-volcanic rocks interaction and origin of hydrothermal solution at midoceanic ridges suggest that Kuroko ore fluids can be produced dominantly by seawater-volcanic rock interaction.
1.4. Epithermal vein-type deposits Epithermal vein-type deposits can be divided into four types based on total metal produced and metal ratio: base-metal type, precious-metal (Au, Ag) type, Sb-type and Hg-
84
Chapter 1
type based on the association of metals. Table 1.2 summarizes estimated total productions of Au and Ag, and Ag/Au. Ag/Au of base-metal type is higher (58 by weight ratio) than that of Au-Ag type (13 by weight ratio). The characteristic features of base-metal and precious-metal types are summarized below and those of Sb- and Hg-type deposits are described in section 1.7.
1.4.1. Geological characteristics 1.4.1.1. Distribution Epithermal Au-Ag vein-type deposits occur widely in the Japanese Islands (e.g., Northeast Hokkaido, Southwest Hokkaido, Northeast Honshu, Sado Island, Izu Peninsula, Kyushu) (Fig. 1.62). The deposits occur in young (late Miocene, and Pliocene), subaerial volcanic regions and in submarine-altered volcanic regions (so-called Green tuff region). The distribution area is located at arc-arc junction (Kubota, 1994) and margin of Green tuff region (Figs. 1.5 and 1.6). These deposits are the principal gold and silver producers in Japan (Table 1.2). Epithermal base-metal vein-type deposits are distributed in the Green tuff region (Southwest Hokkaido, Northeast Honshu) (Fig. 1.62). The distribution area of this type of deposits is nearly same as that of Kuroko deposits. For example, large deposits (Osarizawa Cu-(Au); Ani Cu-Au; Hosokura Pb-Zn deposits) occur in Northeast Honshu, but are more widely distributed in the Green tuff region than Kuroko deposits. The base-metal vein-type deposits in Northeast Japan occur chiefly in Oligocene and early-middle Miocene submarine strata in the members of the Monzen (60-25 Ma), the Daijima (25-15 Ma), and the Nishikurosawa (16-14 Ma), but sometimes they are found in those of the Onnagawa (13-7 Ma) and the Funakawa (7-2.5 Ma) stages of late Miocene (Fig. 1.63) (Nakamura and Hunahashi, 1970). Their distributions are mainly confined to the marginal part of the depressional sedimentary basin which formed during the Nishikurosawa-Onnagawa stage, and are characterized by the presence of rhyolite and dacite or the presence of the so-called Tertiary granite (Nakamura and Hunahashi, 1970). In Southwest Hokkaido, Cu-Pb-Zn veins and Kuroko deposits are distributed in the centre of a volcanic depression zone where submarine rocks predominate, while Au-Ag veins do not (Shikazono and Shimizu, 1988a). Many Cu-Pb-Zn vein-type deposits are hosted by organic sedimentary rocks such as shale and mudstone but almost all Au-Ag deposits occur in altered volcanic rocks. This difference in the host rocks affects the chemical features of ore fluids (fo2, fs2, fco2) (section 1.4.4). 1.4.1.2. Age of mineralization The age of formation of epithermal vein-type deposits can be estimated from K-Ar ages of K-bearing minerals (adularia, sericite) in veins and in hydrothermal alteration zones nearby the veins. A large number of K-Ar age data have been accumulated since the work by Yamaoka and Ueda (1974) who reported K-At age data on adularia from Seigoshi Au-Ag (3.7 Ma) and Takadama Au-Ag deposits (8.4 Ma). Before their publication on the K-Ar ages of these deposits it was generally accepted that epithermal
85
Miocene-Pliocene Hydrothermal Ore Deposits
Sanru II
I '
Hokuryu
.KonomQi
I
0 I00 200km
Q
Todor¢
Yakumo J okok u
Imaiishizaki
Osorizawa
Jlwoto 4o °
Takachi
11nnai iYatanl IHanda Koruizowa
Mizobe
Takeno
Hoshin~
Omidani tAkc
Bojo
\
~1~Seigoshi ..----Yugashimo --Nowaji .Kawozu
35 o
"~ T ~ Toi :ai~1 L ~Sakoshiodomori
1J
Okuchi Hishikari mgano
2
J
~ushikino
lourQ
iwoto I 133 °
I 138°
Figure 1.62. Location of epithermai-type deposits in Japan (Shikazono and Shimizu, 1988a). l: Green tuff and subaeriaI volcanic region of Tertiary/Quaternary ages, 2: Main PaIeozoic/Mesozoic sedimentary terranes, 3: Main metamorphic terranes. TTL: Tanakura tectonic line, ISTL: Itoigawa-Shizuoka tectonic line, MTL: Median tectonic line. Open circle: epithermaI Au-Ag vein-type deposits, solid circle: epithermal base metal vein-type deposits, open triangle: epithermal Au disseminated-type deposits.
vein-type deposits in Tertiary volcanic zone formed in Miocene age, the same as Kuroko deposits because Miocene rocks host the veins. The K - A r age data are summarized in Figs. 1.64 and 1.65. It is obvious in these figures that (1) ages of formation of epithermal vein-type deposits vary widely from 15 to 1 Ma, but are mostly 6-1 Ma, (2) epithermal vein-type deposits have been formed
Chapter I
86
/
Figure 1.63. Distribution of vein-type deposits in a part of the inner belt of Northeast Japan. A: the sedimentary members of the Onnagawa stage, B: the sedimentary members of the Nishikurosawa stage (Nakamura and Hunahashi, 1970).
[] ° _
>,
Au-Ag Vein
[ ] Cu-Pb-Zn Vein
Kuroko
t-
< "5 E Z
15
10
5
1
Age (Ma) Figure 1.64. Ages of formation of Neogenie base-metal vein-type, Au-Ag vein-type and Kuroko deposits, estimated from K-Ar age and paleontologic data (Shikazono, 1987b).
later than Kuroko deposits (15-16 Ma), (3) base-metal veins have been formed earlier than Au-Ag veins (Sawai and Itaya, 1996) (Fig. 1.65), and (4) the frequency of ages of mineralization has three peaks: 10 Ma, 5-6 Ma, and 2-1 Ma.
Miocene-Pliocene Hydrothermal Ore Deposits
87
Au-Ag vein
I
I
I
r
,
Cu-Pb-Zn vein
°+f +I 0L
Mn vein
F I~
Mn strata-bound
z , ,, Barite velnF-- ]
F-'l
Massive barite Pyrite vein
L,
II,
,,
Kuroko
K-At age(Ma) 10
L 1=
Figure 1.65. Histogram of K-Ar ages of hydrothermal ore deposits in the Shakotan-Shikotsu district, Hokkaido (Sawai and Itaya, 1996).
1.4.1.3. Volcanic activity related to mineralization Volcanic rocks associated with this type of mineralization are mostly andesitic rocks. For example, Watanabe (1990b) found that the age of andesitic rocks (flat lava) becomes younger eastward, correlating to the age of epithermal mineralization in Southwest Hokkaido, suggesting a genetic relationship between andesitic volcanism and mineralization. Izawa and Urashima (1989) showed a good correlation of ages of andesitic volcanic rocks and A u - A g deposits in south Kyushu. Their ranges are 4 Ma-1 Ma. However, west of Izu Peninsula, this correlation is unclear and bimodal volcanism (acidic and basaltic activities) seems related to the epithermal A u - A g mineralization.
88
Chapter 1
1.4.1.4. Metal enriched and metal ratios The ore deposits can be classed into two types based on the types of associated metals: Au-Ag rich deposits (Type A) from which Au and Ag are produced as main products, and base metal (Cu, Pb, Zn, Mn, (Sn), (W), (Bi), (Mo), (Sb)) rich deposits (Type B) from which Au and Ag are recovered as byproducts. The deposits are associated with felsic and intermediate volcanic rocks but generally not with felsic plutonic rocks. In Japan Au-Ag deposits associated with granitic rocks (e.g., Au-Ag vein-type deposits in Kitakami) occur commonly. However, these plutonic-type deposits are not described here. Total tonnages of production of Au, Ag and other associated base metal elements were estimated from various records (e.g., Geological Survey of Japan, 1980), as shown in Table 1.2. The Ag/Au ratio of the Type 1-A is lower than that of the Type 1-B from which base metals other than Au and Ag have been produced. In general, the vein-type deposits which have produced large amounts of the base metal elements especially Pb and Zn produce small amounts of Au, but sometimes large amounts of Ag (e.g., Toyoha, Yagumo, Hosokura, Yatani). The vein-type deposits which are rich in Mn have also high Ag/Au (e.g., Ohe-Inakuraishi, Yagumo, Rendaiji). Some vein-type deposits which have produced Cu tend to have large Ag and Au amounts (e.g., Furokura, Osarizawa, Kishu, Ikuno, Tada), but some small deposits which have produced small amounts of Cu have low Ag/Au ratios (e.g., Chitose, Takanosu, and Isobe-Koyama). These deposits, except Chitose from which very small amount of Cu has been produced, are stockwork in form and not typical vein-type deposits. Based on the total production data on individual mines, total production of Au and Ag and Ag/Au ratio from each type of deposits were calculated (Table 1.2). Type 1-A deposits have produced the largest amount of Au and Ag. Total production of Au from the Type 1-B and 2 is small. Ag production from the Type 1-B is lower than but not so different from that of Type 1-A. Average Ag/Au ratio of Type I-A is about 13 and that of Type 1-B is 58.
1.4.2. Mineralogical characteristics 1.4.2.1. Metal zoning Orebody zoning (Park and Macdiarmid, 1963) is observed in Cu-Pb-Zn deposits. For example, in Osarizawa deposit, which is one of the largest Cu-Pb-Zn deposits in Japan, ore metal zoning from deeper to shallower parts is Cu -+ ZnPb --+ AuAg. In the Yatani deposits, Pb-Zn ore occurs in the deeper part, while Au-Ag ore in the shallower part (Figs. 1.66 and 1.67). However, it is uncertain that Au-Ag vein changes continuously to Pb-Zn vein in deeper parts. Orebody zoning is not observed in Au-Ag veins, although the Au/Ag ratio of ore changes considerably with depth. District and regional zonings (Park and Macdiarmid, 1963) are generally not found in Cu-Pb-Zn mine district nor in Au-Ag mine district. 1.4.2.2. Ore minerals Epithermal base-metal vein-type deposits are characterized by the abundant occurrence of sulfides (chalcopyrite, pyrite, sphalerite, galena), and a scarcity of Au-
Miocene-Pliocene Hydrothermal Ore Deposits
89
Au.AgVein
Fault
/" J
$1IJ
/ ~o H 2
%." "'dz.., ,." , ~ ' - -
--'-¢~-'~----
A Bo
wTu o
t
I
~' WH "
500 m
_~..~P K3 ,K2 - ' ~ - . .
!
- ' " '--Esu ~ ~ -
" eo
\
B Figure 1.66. Distribution of Au-Ag veins and Zn-Pb veins of the Yatani deposits. ETI: East-Tengu No. 1, ET2: East-Tengu No. 2, TI: Tengu No. 1, T2: Tengu No. 2, T3: Tengu No. 3, KH: Kanizawa-Honpi, WT: West-Tengu, WT2: West-Tengu No. 2, WH: West-Honpi, H2: Honpi No. 2, KU: Kanizawa Uwabanhi, YH: Yatani-Honpi, W7U: W7-Uwabanhi, E8U: E8-Uwabanhi, E3U: E3-Uwabanhi, N01F: No. 1 fault, N02F: No. 2 Fault (Shikazono and Shimazaki, 1985).
minerals. Ag-minerals (e.g., argentite, pyrargyrite, polybasite) are commonly recognized. Pyrrhotite, magnetite, hematite, and marcasite are occasionally observed. Electron microprobe analyses have revealed that many varieties of Ag mineral occur in the base metal and Au-Ag deposits. The most important features on the occurrences of Au-Ag minerals can be summarized as follows. Au-Ag vein-type deposits (Type l-A): (1) Se minerals such as naumannite, aguilarite, Se-bearing argenitie, and Se-bearing polybasite are found in some deposits (e.g. Sanru, Koryu, Takadama, Kushikino). (2) Au-Ag-Te minerals are rare throughout all of the studied deposits, but sometimes found in some deposits (e.g., hessite from Fuke, Ohkuchi, Arakawa). (3) The common Ag-Au minerals are argentite, As-Se minerals and Ag sulfosalts (polybasite, pyrargyrite, pearceite). (4) Bi, Pb, Zn, and Sn-bearing Ag minerals have not been found. (5) Electrum is abundant, compared with Type 1-B deposits. Native silver is poor in amounts. Base metal vein-type deposits (Type l-B): (1) The common Ag minerals are argentite, and Ag sulfosalts (pyrargyrite, polybasite). Ag sulfosalts are abundant in the late stage of mineralization and argentite occurs in the early stage of mineralization (e.g., Ohe-Inakuraishi, Toyoha).
Chapter 1
90
NO3F ¢,NO2F Ac
~B
, ~/i/~ /
K3
K2 i
?:
i
"/7 •
1
11 :: WH
, /~w7
,/ I
I I
o I
I
200 m 1
Figure 1.67. Vertical cross section along A-B in Fig. 1.66. Abbreviations are the same as those in Fig. 1.66 (Shikazono and Shimazaki, 1985).
(2) Se and Te minerals have not been found in these deposits. (3) Bi, Pb, Zn, and Sn-bearing Ag minerals are rarely found. (4) Abundance of electrum is small, although native silver is abundant in some deposits (e.g., Toyoha, Ikuno). A large number of analytical data on chemical composition of sphalerite are available (Shikazono, 1974a; Watanabe and Soeda, 1981 ). The FeS content of sphalerite from epithermal base-metal vein-type deposits varies widely mostly from 1 to 20 mol% (Fig. 1.68). The FeS content of sphalerite from epithermal Au-Ag vein-type deposits is low, mostly less than 1 mol% (Fig. 1.68). The FeS content of sphalerite from epithermal Se-type is slightly higher than that of epithermal Te-type. The FeS content of sphalerite depends on iron minerals coexisting with sphalerite. The FeS content of sphalerite coexisting with hematite is low (0.5-3.0 mol%), while that of sphalerite coexisting with pyrrhotite is high (mostly 10-20 mol%; Shikazono, 1975). In individual deposits (Toyoha Pb-Zn, Hosokura Pb-Zn deposits), the FeS content of sphalerite coexisting with pyrite varies widely in a range of 1-15 mot%. Mn and Cd contents of sphalerite from epithermal base-metal vein-type deposits are low except Mn content of sphalerite coexisting with alabandite which contains 4.4 wt% Mn (maximum value) (Shikazono, 1975).
91
Miocene-Pliocene Hydrothermal Ore Deposits
Frequency
Frequency Kuroko-typedeposits
EpithermalAu-Ag vein-typedeposits
m 10
15
20 FeS mole %
1"0
1'5
20 FeS mole °/~
Frequency
~rL,~ EpithermalCu-Pb-Zn
•
5
10
15
20 FeS mole%
Figure 1.68. Iron content of sphalerite from Kuroko, epithermaI Au-Ag vein-typeand epithermal base metal vein-typedeposits (Shikazono, 1977a). The Ag content of electrum from epithermal Au-Ag vein-type deposits is mostly in a range of 40-70 atomic% (Fig. 1.69). The Ag content of electrum from the Se-type varies widely (Fig. 1.70). The average Ag atomic% is 50 to 55. The Ag content of electrum from the deposits associated with Te mineralization (e.g., Chitose, Fuke, Takeno) is low, ranging from 26.0 to 40.6 atomic% (Fig. 1.70). Very few data on the chemical composition of electrum from epithermal basemetal vein-type deposits are available. However, it is evident that the Ag content varies widely (Fig. 1.71). The Ag content of electrum from the Osarizawa and Okuyama Cu deposits is low (NAg (Ag atomic%) = 8.6-17.7), while the Ag content of electrum from Pb-Zn-Mn deposits (Toyoha, Oe, Inakuraishi, and Imai-Ishizaki) is high (NAg = 60-80). Motomura (1986) reported that the Ag content of electrum from these deposits is higher than that from epithermal Au-Ag vein-type deposits. The geochemical implication of the Ag content of electrum is discussed in section 1.4.4. Chemical compositions of tetrahedrite-tennantite from epithermal base-metal vein-type deposits are characterized by (1) wide compositional variations, and (2) higher Zn and Sb contents and Ag and lower Fe, As, and Cu contents, compared with Kuroko deposits (Shikazono and Kouda, 1979).
92
Chapter I
100
0 ¢OLL
I
5O
. L• 0
20
40
60
80
100
NAg Figure 1.69. Frequency histogram for the Ag content of electrum from epithermal Au-Ag vein-type deposits in Japan (Shikazono and Shimizu, 1988a).
The differences in Zn/Fe ratio of tetrahedrite-tennantite in epithermal vein-type and Kuroko deposits and that of sphalerite in these deposits can be interpreted in terms of the following exchange reaction: ((Cu, Ag)10Zn2(As,Sb)4S 13)tet "b (FeS)sph = ((Cu, Ag)10Fe2(As,Sb)4S 13)tet Jr- (ZnS)sph (1-21) where (FeS)sph, (ZnS)sph, are FeS and ZnS components in sphaterite, and ((Cu, Ag)]0 Fe2(As, Sb)4Sl3)tet and ((Cu, Ag) ioZn2(As, Sb)4S13)tet components in tetrahedrite-tennantite, respectively.
Miocene-Pliocene Hydrothermal Ore Deposits Te t y p e d e p o s i t s
• n=5
Date Okuyama Kushikino Agawa Chugu Fuke Okuchi Chitose Takeno Sado Kato
93
.,~n=lO • n=2
--~=n=15 = =-n=69 ~n=lO *n=17 -- = •n=6 -- --= n = l • n=l
Se t y p e d e p o s i t s Sakoshi Hishikari Koryu Chitose Sanru Kushikino Takadama Yatani Nebazawa Omidani
, n=61
u
n=20,~ •
L
L n=22 = n=41--
=
• n=8 =
=
= = n=44 i
I
i
I
I
0
I
I
I
50
I
I
NAg
I
100
Figure 1.70. Ag content (atomic fraction of Ag) of electrum from the Te-type arid Se-type deposits (Shikazono et al., 1990). n: number of analyses.
30-
>20o c-" O" LL
10"
0
2"0
40
60
80
100 NAg
Figure 1.71. Frequency histogram for the Ag content of electrum from epithermai base metal vein-type deposits in Japan (Shikazono and Shimizu, 1988a).
If equilibrium is attained for the above reaction, equilibrium constant (K1-21) is expressed as, KI-2I =
(azntet/aFetet)/(aZnspJaFesph)
(1-22)
94
Chapter 1
where aZn~et, aFetet, aZnsph, and aFesph are activities of tetrahedrite-tennantite and activities of ZnS and FeS in sphalerite, respectively. This equation suggests that the Fe content of tetrahedrite-tennantite positively correlates with that of sphalerite at constant temperature and pressure, indicating Fe and Zn contends of tetrabedrite-tennantite are useful to estimate physicochemical parameters (fs2, fo2 etc.) as well as Fe content of sphalerite, although detailed study on thermochemical properties of tennantite-tetrahedrite solid solution is still needed. The chemical compositions of Ag-minerals have been obtained from several A u Ag and base-metal vein-type deposits (Shikazono, 1978b; Ohta, 1991, 1992). For instance, Shikazono (1978b) found that argentite (acanthite) from epithermal A u - A g vein-type deposits (Seigoshi, Ohmidani) contains appreciable amounts of selenium but that from epithermal base-metal vein-type deposits (Toyoha, Ikuno) does not. High selenium contents of galena and Ag sulfosalts (polybasite, pyrargyrite) from epithermal A u - A g vein-type deposits (Kushikino) were reported by Takeuchi and Shikazono (1984) (Table 1.11). The geochemical implications of selenium contents of sulfides for the physicochemical environment of epithermal ore deposition will be discussed in section 1.4.4.
1.4.2.3. Gangue minerals Quartz is the most abundant gangue mineral. It occurs commonly in Au-Ag and P b - Z n deposits but is scarce in Cu deposits. Chalcedonic quartz coexisting with Au-Ag minerals occurs abundantly in A u - A g deposits. Amethyst is generally rare and occurs as a late-stage mineral in Au-Ag and Pb-Zn deposits. Magnetite is common in P b - Z n - M n and Cu deposits but has not been reported in A u - A g deposits. It commonly coexists with other iron minerals such as hematite, pyrite, pyrrhotite, siderite, and chlorite and also occurs in both the main stage of sulfide mineralization and in the late stage of mineralization. The occurrence of hematite is generally more widespread than magnetite in all types of deposits, especially Cu-Au deposits. It tends to occur in late-stage mineralization, generally later than the sulfide mineralization. Inesite is common in A u - A g deposits, especially in the Izu peninsula (Kato, 1928). Generally it is found associated with highgrade gold-silver ores, as, for example, in the Kakehashi vein of the Kawazu mine. It also occurs rarely in Pb-Zn deposits such as Toyoha and Tatsumata. The occurrence of other M n - C a silicates such as johannsenite, bustamite, rhodonite, pyroxmangite, tephroite, and penwithite has been reported from A u - A g and Pb-Zn deposits, but these minerals are not common. They have not been reported from Cu deposits. Hydrated calcium silicate minerals such as xonotlite, truscottite, and gyrolite are rare but have been reported from several Au-Ag deposits. They do not coexist with A u Ag minerals but instead are found with quartz, carbonates, and johannsenite. However, in the Keisen No. 3-2 vein in the Hishikari Au-Ag deposits, a close association of electrum with truscottite, smectite and calcite is observed (Imai and Uto, 2001). Prehnite is found with A u - A g minerals in the A u - A g vein (Kanisawa vein) of the Yatani P b - Z n - A u - A g deposit but is not found in the Pb-Zn vein (Yatani-Honpi vein).
Miocene-Pliocene Hydrothermal Ore Deposits
95
Calcium silicates such as wairakite, epidote, prehnite, laumontite, and stilbite are common in the wall rocks of some A u - A g deposits in the Izu peninsula. Epidote occurs as a gangue mineral coexisting with sulfides and quartz in some Cu deposits, but none of the other above-mentioned Ca and Mn silicates have been reported from these deposits. Laumontite is a common mineral in propylite, which is the host rock for A u - A g deposits. Other zeolites such as mordenite and dachiardite are not generally common, but they are the main gangue minerals associated with A u - A g minerals in the Ohnoyama and Awagano A u - A g deposits. Chlorite is abundant in Cu-Pb-Zn-rich deposits but is scarce in Au-Ag-rich deposits. Fe chlorite is the most common and Fe-Mg chlorite is subordinate (Shirozu, 1969). Almost all of the chlorite is classified as orthochlorite which can be regarded as part of the clinochlore~taphnite solid solution series. In general, chlorite is intimately associated with sulfide minerals such as sphalerite, galena, pyrite, chalcopyrite, and pyrrhotite. A 7 A septechlorite was reported from the Toyoha Pb-Zn deposits (Sawai, 1980). Interstratified chlorite-smectite and vermiculite-saponite are rather common minerals in A u - A g deposits (e.g., Yoneda and Watanabe, 1981), but they have not yet been reported from other deposits. Smectite commonly occurs in A u - A g deposits, and to a lesser extent in Pb-Zn deposits, but only rarely in Cu deposits. The occurrence of beidelite in the Seigoshi and Ohkuchi A u - A g deposits has been described (Nagasawa et al., 1981), and sericite is a common gangue and alteration mineral in C u - P b - Z n rich deposits, where it occurs as a late-stage mineral (Nagasawa et al., 1976). Sericites (Watanabe et al., 1982) from A u - A g deposits occur in the wall rocks as alteration products. In addition, kaolinite commonly occurs as a late-stage mineral in some A u - A g deposits, but not together with the ore minerals. Sometimes kaolinite has been found as an alteration product of adularia after main-stage mineralization. Carbonate minerals occur in almost all the vein-type deposits. In general, they are more abundant than Ca and Mn silicates, but their abundance varies widely with different types of deposits. Large amounts of Mn carbonates (rhodochrosite and manganoan calcite) occur in P b - Z n - M n deposits, moderate amounts in Pb-Zn and Cu deposits, and small amounts in A u - A g deposits. Calcite is abundant in all types of deposits. Siderite is common in Cu deposits, especially in C u - A u deposits, but it is uncommon in A u - A g and Pb-Zn deposits. Siderite coexists with hematite, pyrite, chlorite, and rarely with magnetite; it is considered to have been precipitated after the main stage of sulfide mineralization. Other carbonates such as ankerite, dolomite, and kutnahorite are not common. Ankerite has been reported from several C u - P b - Z n deposits where it is found with sulfides and also with oxides such as magnetite and hematite without sulfides. Dolomite is found in some A u - A g deposits but is rare in other deposit types. Siderite and ankerite coexist with oxides and sulfides, but calcite generally occurs as a late-stage mineral associated with quartz and sericite. Mn carbonates are found with sulfides in P b - Z n - M n deposits such as Ohe and Inakuraishi, but this assemblage is not common. Carbonates from epithermal base-metal and A u - A g vein-type deposits are different. Mn- and Fe-carbonates are common in base-metal vein-type deposits and calcite is abundant in A u - A g vein-type deposits. Shikazono (1973) revealed that the iron content
T A B L E 1.1 l R e p r e s e n t a t i v e analytical data on ore m i n e r a l s from K u s h i k i n o A u - A g d e p o s i t (Takeuchi and S h i k a z o n o , 1984) S a m p l e No.
Period
Cu
Au
Aq
Zn
Fe
Cd
Mn
Sb
As
S
Se
Total
FeS a (tool %)
Electrum 77080212e, L6
IV
-
65.0
33.0
.
76081806a, L6
lI
-
60.5
38.7
.
.
.
.
.
7 7 0 7 2 6 0 3 a , 1.8
lib
-
60.0
38.5
.
.
77072601a, L8
lla
-
63.0
35.0
.
76081810c, L9
lib
-
57.0
42.0
.
.
.
.
.
.
.
.
99.0
76081809f, L9
llb
-
67.0
34.0
.
.
.
.
.
.
.
.
101.0
Za, L 6
IV
2.63
-
69.5 ~
0.09
0.00
-
-
6.49
2,22
11.84
6.36
99.14
77080212b, L6
IV
3.69
-
67.57
0.(X/
0.00
-
-
8.62
0.74
11.49
6.91
99.02
78031601 c, L B
Ila
3.09
-
69.63
0.00
0.00
-
-
9.44
(t.79
14.81
1.66
99.42
77072602f, L8
Ila
2.67
-
67.69
0.01
0.00
-
-
11.24
0.33
14.32
3.53
99,79
77072602i, L8
Ma
3.01
-
66.24
0.00
0.00
-
-
10.60
0.75
14.90
3.75
99.25
77072602d, L8
lla
3,t8
-
69.52
0.(X)
0.00
-
-
10.05
0,93
15.09
1.20
99,97
7 7 0 7 2 9 0 l c , 1.8
I1
2.77
-
68.32
0.11
0.00
-
-
9,98
0.55
12.31
5.57
99.61
7 7 0 7 2 9 0 B c , 1.8
llb
2.48
-
68,15
0.02
0.00
-
-
9.90
0.25
ll.90
6.71
99.41
7 7 0 7 2 7 0 1 a , 1.8
lib
2.68
-
68.91
0.(X)
0.00
-
-
9.65
0.51
12.26
5.92
99.93
7 6 0 8 1 8 0 9 c . L9
IIb
3.80
-
65.78
0.10
0.31
-
-
9.85
0.37
12.04
7.30
99,55
7 6 0 8 2 3 0 3 c , 1.9
IIb
3.84
-
70.83
0.10
0.00
-
-
9.14
0.36
12.86
2.99
100.12
76081809b, L9
IIb
4.12
-
65.09
0.25
0.63
-
-
9.84
0.54
11.50
7.57
99.52
76081806a, L9
IIb
4.13
-
71.36
0.44
0.26
-
-
7.35
0.46
10.01
5.01
99.02
76081806c, L9
Ilb
4.32
-
68.50
0.25
0.00
-
-
9.18
0.44
12.57
4.68
99.94
. . .
.
.
.
. .
. .
. . .
.
. . .
.
98.t) .
99.2
.
98.5
.
98.0
Polybasite
¢5
T A B L E 1.11 ( c o n t i n u e d ) I
Sample No.
Period
Cu
Au
Aq
Zn
Fe
Cd
Mn
Sb
As
S
Se
Total
FeS ~
~.
(tool % ) Tetrahedrite
77080221a, L6
II
23.68
-
20.20
5.24
1.29
-
26.19
1.25
22.30
0.01
77080212e, L6
IV
23.92
-
19.16
4.81
1.71
-
-
27.79
0,02
22.60
0.134
100.05
77072901c,
18
IIb
23.01
-
20.60
4.34
2.02
-
-
26.07
1.16
23.00
I).06
100.26
77072912c, L8
lib
32.42
-
7.14
5,97
1.05
-
-
27.23
1.71
24.00
0.00
99.52
76081809b, L9
IIb
22.99
-
20.90
3.86
2.57
-
26.07
1.23
23.37
0,25
101.24
76081B06b,
1.9
IIb
28.54
-
12.10
6,48
0.57
-
-
27.95
0.73
23.54
0.02
99.93
76081809a,
1.9
IIb
22.02
-
211.39
3,87
2.64
-
-
26.33
1.34
22.88
0.00
100.47
100.16 ~-~
~z Naumanni~
77080212,
1.6
7 6 0 8 2 3 0 3 , 19
IV
-
-
74.17
-
0.64
25.69
100.50
lib
-
-
80.57
-
8.96
7.65
97.18
100.45
~"
Pyrargy~te
1, L d
Ilb
0.00
-
59.98
0.00
0.00
77072602A, LB
lla
0.00
-
59.82
0.00
0.00
77072602B, L8
IIa
0.00
-
60.53
0.00
0.00
_
m
63.84
m
m
65.13
-
-
22.46
0.37
16.45
1.19
-
22.06
0.56
17.08
0.00
99.52
-
-
22.42
0.15
17.12
0,26
100.48
0.76
0.78
0.05
-
-
32.59
-
0.06
0.00
0.10
-
-
33.11
Sphalerite
77072901a,
18
lib
0.10
77072909e, L8
IV
0.00
77072601e, L8
IIa
0.10
62.72
1.56
0.52
0.00
-
-
33.56
76081801a, L9
lib
0.24
62,93
1.38
0.42
0.111
-
-
33.86
76081802b, L9
IIb
0.18
63.65
1.80
0.39
0.00
-
-
7 6 0 8 1 B 10b, L 9
IIb
0.33
63.29
1.98
0.83
0.00
-
76081810c,
lib
1.73
62,14
2.65
0.73
0.01
-
19
r
98.12
1.2
99.40
0.1
-
98.46
2.5
-
98.84
2.0
33.67
-
99.69
2.8
-
33.08
-
99.51
3.0
-
32.40
-
99.66
2.1
a F e S : C o r r e c t e d v a l u e e x c l u d i n g the e f f e c t o f X - r a y s c a t t e r i n g c a u s e d b y c h a l c o p y r i t e .
"--3
98
Chapter 1
(FeCO3) of Mn-carbonates varies widely in a range of 10-3-10 -1 mole fraction. Siderite from Au-Ag vein-type deposits (Ohmori) contains appreciable amounts of zinc (0.8-5.8 wt% as ZnO) (Shikazono, 1977b). Adularia is abundant in Au-Ag deposits, where it is commonly found with AuAg minerals; only rarely does it occur in Pb-Zn and Cu deposits. Albite is very rare and is reported only from the Nebazawa Au-Ag deposits. Barite is a common gangue constituent in P b - Z n - M n deposits, especially those in the southwestern part of Hokkaido and the northern part of Honshu, where it is usually a late-stage mineral coexisting with carbonate and quartz but rarely with sulfide minerals. Other rare gangue minerals include fluorite, apatite, gypsum, bementite, rutile, and sphene, but they have not been studied. Main gangue minerals of the Se-type deposits comprise quartz, adularia, illite/ smectite interstratified mixed layer clay mineral, chlorite/smectite interstratified mixed layer clay mineral, smectite, calcite, Mn-carbonates, manganoan calcite, rhodochrosite, Mn-silicates (inesite, johannsenite) and Ca-silicates (xonotlite, truscottite). In comparison, the Te-type deposits contain fine-grained quartz, chalcedonic quartz, sericite, barite, adularia, chlorite/smectite interstratified mixed layer clay mineral and rarely anatase. Carbonates and Mn-minerals are very poor in the Te-type deposits and they do not coexist with Te-minerals. Carbonates are abundant and barite is absent in the Se-type deposits. The grain size of quartz in the Te-type deposits is very fine, while large quartz crystals are common in the Se-type deposits although they formed in a late stage and do not coexist with Au-Ag minerals. Principal gangue minerals in base-metal vein-type deposits are quartz, chlorite, Mn-carbonates, calcite, siderite and sericite (Shikazono, 1985b). Barite is sometimes found. K-feldspar, Mn-silicates, interstratified mixed layer clay minerals (chlorite/smectite, sericite/smectite) are absent. Vuggy, comb, cockade, banding and brecciated textures are commonly observed in these veins. The predominant gangue minerals vary with different types of ore deposits; quartz, chalcedonic quartz, adularia, calcite, smectite, interstratified mica/smectite, interstratified chlorite/smectite, sericite, zeolites and kaolinite in Au-Ag rich deposits; chlorite, quartz, sericite, carbonates (calcite, rhodochrosite, siderite), and rare magnetite in Pb-Zn rich deposits; chlorite, sericite, siderite, hematite, magnetite and rare epidote in Cu-rich deposits (Sudo, 1954; Nagasawa et al., 1976; Shikazono, 1985b).
1.4.2.4. Hydrothermal alteration zoning Among the epithermal vein-type deposits in Japan, four major types of hydrothermal alteration can be discriminated. They are: (1) propylitic alteration, (2) potassic alteration, (3) intermediate argillic alteration, and (4) advanced argillic alteration. The definitions of these types of alteration are mainly based on Meyer and Hemley (1967) and Rose and Burt (1979) who classified the hydrothermal alteration in terms of alteration mineral assemblages. Representative propylitic alteration minerals include epidote, albite, carbonates, quartz, chlorite, sericite, and smectite. The less common minerals are mixed-layer clay minerals such as chlorite/smectite and sericite/smectite and zeolite minerals. The term "propylite" is widely used to describe altered volcanic rocks recognized
Miocene-Pliocene Hydrothermal Ore Deposits
99
in mine areas. However, the term is somehow ambiguous, because it is defined differently by different investigators (e.g., von Richthofen, 1886; Kato, 1928). Propylite is defined here as the andesitic, dacitic and rhyolitic rocks which have undergone the alteration to form the characteristic minerals such as chlorite, epidote, albite, K-feldspar, etc. Probably there are two different origins of this type of alteration. The propylite may be formed either by metamophism or by hydrothermal alteration. However, it is extremely difficult to distinguish rigidly between these two origins. Here the studies on the propylitic alteration which is considered to be intimately related to igneous activity in the mine area are summarized. Usually lateral and vertical zonations are observed in the area of propylitic alteration. Generally epidote, actinolite and chlorite tend to occur in the central and deeper parts, while at the marginal and shallower parts zeolites, mixed-layer clay minerals and smectite are commonly observed. The following zonation is generally recognized from the central to marginal parts: chlorite --+ chlorite/smectite -+ smectite. This zoning pattern is observed in the Fuke Au-Ag district (Inome et al., 1981), Seigoshi Au-Ag district (Shikazono, 1985a), Hosokura Pb-Zn district (Suzuki et al., 1982), Toyoha PbZn district (Yoshitani, 1971; Sawai, 1984), Miyatamata, Arakawa and Nissho C u - P b Zn districts (cited in Utada, 1980). At the central parts of some mine districts listed above, intrusive plutonic bodies have been observed. The formations of high temperature alteration minerals including epidote, actinolite, prehnite and wairakite as suggested by Yoshitani (1971), Fujii (1976), and Shikazono (1985a) are considered to be attributed to these intrusive activities. Zeolite minerals (wairakite, laumontite etc.), mixed-layer clay minerals and smecite occur in the upper part of the propylitically altered rocks (e.g., Seigoshi, Fuke, Kushikino), but they are sometimes poor in amounts. Generally carbonates are more abundant in the mine area as in the Toyoha district. Temporal relationship between the formation of high temperature propylitic alteration minerals (epidote, actinolite, prehnite) and low temperature propylitic alteration minerals) (wairakite, laumontite, chlorite/smectite, smectite) in these areas (Seigoshi, Fuke, Kushikino) is uncertain. Potassic alteration has been recognized in the Numanoue (Otagaki, 1951), Konomai (Urashima, 1953), Takarakura (Utada, 1980), Seigoshi (Shikazono and Aoki, 1981; Shikazono, 1985a; Imai, 1986), Hosokura (Suzuki et al., 1982; Konno et al., 1984), Chitose (Takatori and Nohno, 1985), Karuizaiwa (Sugaki et al., 1986) and the Ikutahara district, Hokkaido (e.g., Yahagi, Ryuo) (Matsueda et al., 1992) where adularia occurs dominantly. In the Mikawa (Sudo et al., 1953, Nagasawa, 1961, 1962), Ohe (Tsukada and Uno, 1980), Toyoha (Okabe and Bamba, 1976), Rendaiji (Watanabe and Nagai, 1986) and Kushikino (Izawa et al., 1987), sericitization is dominant. The area of the potassic alteration is not wide, compared with the propylitically altered area. The width of potassic alteration zone away from the vein is generally within several tens of meters (ca. 50 m) (Shikazono and Aoki, 1981; Imai, 1986). The potassic alteration is usually found in the intermediate vicinity of the vein in the epithermal deposits in Japan. Thus it is evident that this type of alteration occurs genetically related to the ore deposition. Lateral zonation from a sericitic envelope to an intermediate argillic envelope is common in the porphyry copper deposits and vein-type deposits in granodioritic rocks
100
Chapter I
(Meyer and Hemley, 1967). However, such lateral and concentric zonation has not been reported from the epithermal vein-type deposits in Japan. Montmorillonite-rich and silicarich zones exist in the upper part of the Au-Ag veins such as the Seigoshi and Takadama (Nagasawa et al., 1981). In contrast to the hardly investigated lateral zonation around Japanese epithermal vein-type deposits, a few examples of vertical zonation are known. Potassic alteration grades upwards into intermediate argillic alteration in the wall rocks for the Toyoha (Okabe and Bamba, 1976), Ohe (Tsukada and Uno, 1980), Chitose (Hasegawa et al., 1981) and Kushikino (Imai, 1986). Advanced arigillic alteration is found at the upper horizon than the sites of potassic and intermediate argillic alterations where the Au-Ag mineralization occurs (e.g., Seigoshi, Yatani, Kushikino, Hishikari). This type of alteration takes blanketform in upper part and vein-form in lower part (Iwao, 1962; Shikazono, 1985a). The conspicuous zonation from upper to lower horizon is known at the Ugusu silica deposit, namely, silica zone, alunite zone, kaolinite zone and montmorillonite zone (Iwao, 1949, 1958, 1962). It is rather difficult to determine the sequence of each type of alteration in a mine area. However, it is widely accepted that the hydrothermal alteration proceeds as follows: propylitic alteration ~ potassic alteration and intermediate argillic alteration -~ advanced argillic alteration. The actual sequence alteration might be more complicated and superimposition of each type of alteration could be common. Usually propylitic alteration precedes the base metal and Au-Ag mineralizations. Potassic and intermediate argillic alterations are nearly contemporaneous with ore deposition. Advanced argillic alteration is also nearly contemporaneous with base-metal and Au-Ag mineralization as found in the Seigoshi and Yatani Au-Ag mine areas based on K-Ar datings (Shikazono, 1985e), although advanced argillic alteration was not caused by the hydrothermal solution responsible for the Au-Ag mineralization. Advanced argillic alteration at the Ugusu silica deposit is inferred to be caused by the mixing of volcanic gas containing SO2 and groundwater (Shikazono, 1985a). Generally, the chemical composition of rocks does not considerably change during the propylitic alteration. The components which are added to the rocks are only H20, CO2 and S (e.g., Okabe and Bamba, 1976). Considerable changes in the chemical composition of rocks occur during the advanced argillic alteration. For example, SiO2 content of highly silicified rocks of the Ugusu silica mine reaches 99% (Iwao, 1962). This silicification is caused by the considerable leaching of elements from the rocks by acid hydrothermal solution except Si and addition of Si from hydrothermal solution.
1.4.2.5. Spatial and geochemical relationships between propylitic alteration and advanced argillic alteration: a case study on the Seigoshi-Ugusu district, central Japan In the Izu Peninsula, located in the middle part of Honshu, more than 20 epithermal Au-Ag vein-type deposits have been mined. Large Au-Ag mines are located in the western part of the peninsula. The Seigoshi mine is the largest one. The country
Miocene-Pliocene Hydrothermal Ore Deposits
101
TABLE 1.12 Generalized stratigraphic succession in the Izu Peninsula and the area surveyed (Shikazono, 1985a)
~ ~e" O° om R" °
' Daruma
andesite lava +500m
iOdoi
andesite lava +500m
~
Tanaba andesite lava .~ =~" Andes te pyroclastics +500m
~'
e- O
K0shimoda andesite lava Andesite volcanic breccia +500m
I-- v
Nekko Dacite
~' "~ ~. R"
dacite lava tuff breccia
+500m I
e= !~
~)
Yagisawa dacite lava andesitic tuff, Formation tuff breccia +1000m •-dacitic tuff u~ breccia andesite lava Tom (~ andesitic tuff, breccia +1000m •= O. Formation sandstone J~ ~ dacitic tuff ~ volcanic conglomerate Ugusu andesite lava pyroclastics >" Formation sandstone conglomerate +500m
rocks in the Seigoshi district have suffered intense propylitic and advanced argillic alterations. The Izu Peninsula is mainly composed of pyroclastic and volcanic rocks of Tertiary-Quaternary age. The general geology of the peninsula has been well studied (Tayama and Niino, 1931), and thus, it is briefly described below. The generalized stratigraphic succession in the Izu Peninsula is shown in Table 1.12. The oldest rock exposed in the peninsula is the Yugashima Group of Miocene age that is composed of submarine andesitic-dacitic pyroclastic and volcanic rocks with small amounts of sandstone. This rock is metamorphosed from zeolite facies to epidoteprehnite-pumpellyite facies. The total thickness of the group is variable from place to place, but, generally exceeds 1500 m. The Shirahama Group which is mainly composed of felsic tuff and sandstone overlies the Yugashima Group conformably or unconformably in different places. The age of this group is considered to be late Miocene to Pliocene. The thickness of this group is also variable from place to place, but the average thickness is more than 1,000 m. These Tertiary rocks are overlain unconformably by subaerial Quaternary andesitic rocks. The generalized geology and schematic stratigraphic succession in the SeigoshiUgusu district is shown in Fig. 1.72. The andesitic and dacitic rocks of Pliocene age unconformably overlie these Tertiary rocks. The Ugusu Formation, corresponding to the lower horizon of the Yugashima group, is characterized by the predominance of andesitic pyroclastic rocks. The maximum thickness of the Ugusu Formation is estimated
Chapter 1
102
l~
....
v v v vVVV'V'V'V
yvVvVvV~vVvVvVvVvVv3yvVvV~vVvVvVvVv vVv ~,Vv v VvVv~v~vV~Vv~v~,,v, v yvV~ v ~WW~VW4 v v v yvV~VvVvV,,Vv~ VvVW,
...........
~AAAA
A A A A A A
A A A A A
~l~
o
0
o o
0 o
o o
0 o
o"o° o
o
0 o
o o
o o
o o
A A ~
Daruma-Odoi Volcano (andesite lava) 1 ~ Tanaba Volcano (andesite lava, pyroclastics) Koshimoda Volcano (andesite lava, volcanic breccia) Nekko Volcano (dacite lava, tuff breccia) Yagisawa formation ........ (dacitic lava, andesitic tuff, tuff breccia dacitic tuff breccia)
[~
Toi formation (andesite lava, andesitic tuff breccia, sandstone, dacitic tuff, volcanic conglomerate) ~ - ~ Ugusu formation (andesite lava, pyroclastics, sandstone, conglomerate) Intrusive rock of diorite porphyry
Figure 1.72. Generalized geological map of the Seigoshi-Ugusuarea (Shikazono, 1985a).
to be more than 500 m. The Toi Formation, corresponding to the upper horizon of the Yugashima Group, conformably overlies the Ugusu Formation. This formation is mainly composed of pyroclastic and volcanic rocks of andesitic composition. The average thickness of this formation is 1,000 m. The Yagisawa Formation, which may correspond to the Shirahama Group, unconformably overlies the Toi Formation. This formation is composed of dacitic and andesitic lavas and pyroclastic rocks. Its total thickness is ca. 1,000 m. Andesitic and dacitic rocks of Pliocene or Pleistocene unconformably overlie the Yugashima and Yagisawa formations. In the northern part of this district, Quaternary andesitic rocks are distributed. These rocks are composed of an alternation of thick lava flows of hypersthene-augite andesite and their autobrecciated parts. Many faults trending generally from north to south are developed, mainly in the Yugashima Group (Fig. 1.73). The axis of the minor foldings in the Yugashima rocks is also generally north to south. Basic and felsic intrusive and dike rocks often occur in the
Miocene-Pliocene Hydrothermal Ore Deposits
t
103
J
,,' ,'~... Advanced Argillic ~" . \ ,' p.. ,' Alteration m T "1 ,' ~Funabara T \ ,'..----0 . . . . 4. area " ~ " : " ~ o ~Seigosi~i'-. " , ~ o i ,• T o"',l ~ ~, ~ "',. . ,,mine . co ~ n e ', ,.~ ,'/ . . , . , Propylltlc 'j .." .,' ', 'J" ,,', ",<~ Alteration .. ,; ~ F9 ,,'~ ,, ~,-" :/Ugusu ,' Mochikoshi /," ,: m ne : \ '. mine
', F / .,a-x" - 2.'. 't
, ;
Amagi mine : , , ' ~ A ' " ?~ ,'
"J'.
" 'J
",3
',.
~, < 3 ,
F
"~
.... -..2~ ', "
5 km
;',. . . . . . . . . . . . ', ~ ' ,
•t
'
'." "
['~2q Diorite-Porphyry ~
~
Silicified Rock of Advanced Argillic Alteration Gold-SilverQuartz Vein Fault
Figure 1.73. Distribution of epithermaI Au-Ag vein-typedeposits, propylitic and advanced argillic alterations and intrusive rocks of diorite prophyry(Shikazono, 1985a).
Yugashima, trending from north to south and from east to west. Occasionally the A u - A g quartz veins are hosted by these intrusive rocks and the upper horizon of the Yugashima Group. These intrusive rocks have suffered propylitic alteration. The host rocks for the advanced argillic alteration are generally younger formations such as Pliocene andesite, although the Yugashima Group rocks have also suffered the advanced alteration occurring as a vein in form. The distribution of the A u - A g vein-type deposits in this district is shown in Fig. 1.73. The propylitic alteration is intimately associoated with these deposits. The vein is composed of rhythmic banding of quartz layers and fine-grained sulfides such as argentite, acanthite, sphalerite, galena, pyrite and chalcopyrite, and electrum. The principal gangue minerals are quartz, calcite, adularia and interstratified chlorite/smectite. Minor minerals are inesite, johansenite, xonotlite and sericite. These gangue minerals except for quartz, adularia, calcite and sericite are not found in the wall rocks. The Seigoshi and Toi deposits occur in the andesitic pyroclastic rocks of the upper horizon of the Yugashima Group and basic intrusive rocks. Distributions of the wallrock alteration minerals from underground in the Seigoshi mine and on the surface near the
Chapter 1
104 West <
> East
',, • Ao~-'. . . . .
...... 2 . . . .
...... \--~ii-Ne~-2-_:-_:-2-_-2yuo
_. . . . ~ , 2 - 1 ..-+
5--4 j
jJ
', ~~ \ "
'
-.
,
,
.:~J/-~,," ~,, opx J
A~
J'~:f-
f~f J'~
•
+-
?I,
o~
\
"
\
'
', ',
I~.
'
/
"
/
T
I
I
I t
e/
',
--
, /l
', ,
#5-, "v ,
.......
~
I
SeigoshiAu-AgVein
Figure 1.74. Zonal sequence of the propylitic alteration in E-W section of the Seigoshi-Toi mine area (Yug = yugawaralite; Heu = heulandite; Stil = stilbite; Opx = orthopyroxene; Mont = montmorillonite;Mor = mordenite; Lm = laumontite;Wr = wairakite; Chl = chlorite; pr = prehnite;ep = epidote; Py = pyrite; Kf = K-feldspar; Cpx = clinopyroxene)(Shikazono, 1985a). Deep(A)
)
Shallow(B)
Mordenite Yugawaralite Heulandite Stilbite Laumontite Wairakite Montmorillonite Chlorite Epidote Prehnite K-feldspar Pyrite
Pyrrhotite Magnetite Sphene
.....
Figure 1.75. Zonal sequenceof the propylitic alteration in section A-B in Fig. 1.74 (Shikazono, 1985a).
mine are shown in Fig. 1.74. This type of alteration has a vertical zonal arrangement (Fig. 1.75). The abundance of the alteration minerals also changes from the portion near the A u - A g - q u a r t z vein to the portion away from the A u - A g - q u a r t z vein. From the deep to shallow and from the portion near the A u - A g - q u a r t z vein to the portion away from the vein, the hydrothermal alteration zoning is observed: (1) epidote-prehnite-K-feldsparchlorite zone; (2) wairakite-laumontite zone; and (3) stilbite-heulandite-smectite zone. The rocks of the underground levels of the Seigoshi mine have suffered alterations (1) and (2). The boundary between these zones cuts the strata with a small angle. The boundary between each zone is gradual. In zone (1), quartz, K-feldspar, epidote, chlorite, prehnite and sphene are predominant alteration minerals. Epidote, prehnite and carbonate replace plagioclase phenocryst. Epidote often occurs as a veinlet with several millimeters wide, together with prehnite. K-feldspar, calcite and quartz tend to occur as a veinlet. Chlorite replaces pyroxene
Miocene-Pliocene Hydrothermal Ore Deposits
105
and also occurs as a veinlet often with pyrite. Orthopyroxene is completely replaced by chlorite, but clinopyroxene is sometimes preserved. Amphibole is often replaced by carbonate. The rim of the original magnetite is replaced by pyrite and sphene. In zone (2), wairakite and laumontite are commonly found. Wairakite occurs as a veinlet together with laumontite. Laumontite occurs as a veinlet and also as filling amygdule. Very small amounts of epidote are found with these zeolite minerals. Small amounts of interstratified sericite/montmorillonite are found. Veinlets of epidote, prehnite, quartz, K-feldspar and chlorite appear at the deeper part of this zone. Yugawaralite occurring as a veinlet is rarely found. In zone (3), stilbite, heulandite, and minor amounts of chabazite and mordenite are found mainly as veinlets and filling amygdule. The rocks of this zone have not been significantly altered. Original mafic minerals such as clinopyroxene and orthopyroxene are altered to smectite but sometimes they are preserved. Small amounts of carbonate minerals are found in this zone. A zonal sequence of opaque minerals is also found. Pyrite is found in zones (1) and (2). Very tiny amounts of pyrrhotite coexisting with pyrite are found in zone (1). Magnetite is common in zone (3) but this mineral is thought to be original. Fluid inclusion studies have been carried out on quartz samples from the veinlets in the Yugashima Group and diorite porphyry and from the A u - A g vein. Homogenization temperatures of inclusion fluids in the quartz coexisting with laumontite and stilbite range widely from ~240°C to 380°C as shown in Fig. 1.76. This wide range of homogenization temperatures and the coexistence of vapor- and liquid-rich fluid inclusions in the same quartz crystal suggest that boiling took place when these zeolites and quartz were precipitated. The homogenization temperatures of fluid inclusions in the quartz which is in contact with epidote and prehnite range from ~235 to 285°C. All fluid inclusions in the quartz are liquid-dominated type and vapor homogenized into liquid. The range of temperatures for each alteration zone can be estimated from the following chemical reactions and thermochemical data available for these reactions laumontite
: wairakite + H20
(1-23)
yugawaralite = laumontite + quartz
(1-24)
yugawaralite = wairakite + quartz ÷ H20
(1-25)
stilbite
= laumontite + 3 quartz + H20
(1-26)
heutandite
= laumontite + 2 quartz + 3 H20
(1-27)
stilbite
= heulandite + 3 quartz + 3 H20
(1-28)
The coexistence of laumontite and wairakite is common in zone (1). If the saturated water vapor pressure is equal to 0.3 of total pressure (Zeng and Liou, 1982), the temperature for equilibrium reaction (1-23) and saturated water vapor pressure are estimated to be ~170°C and 230 bar, respectively (Liou, 1971b). Zeng and Liou (1982) have shown that yugawaralite is stable at less than ~230°C and a total pressure of 500 bar, under the condition of quartz saturation. However, if the activity of SiO2 is not unity, the boundary for reactions (1-24) and (1-25) may shift to lower temperatures. Liou (1971a) studied the equilibrium for reaction (1-26) and showed that the equilibrium
Chapter I
106 o o
0 ¢- "0 ~
°o.
o ¢:.-
v
a
o
o
-8 &
o r-
~0 = <
U~
r'-
"E
=
LZ
2 ~ ~
-800 • 700
O cO
.~0 i~ o -600
"~
n -500
g
-400 • 300 "200
;i . . . . . . ~ ~ . , [ ~ v' :, O ooo,
=~
0
N
;
. .
o:: ooo"
oo%oi ~G e.= p' e,ooo o o 0 0>3 oooo;. :,,. ~~ eoe ee el, ee, ~,ee= q o ~ o e o poo J e 13 (/} O e ~ e e *4e e e e e . . . . .
• loo
o
AdvancedArgiIlic AIteratlon
o i
200 300 400 ( T ' C ) ~ Zeolile Zone EpidoleZone ~
Au-AgVein ,
,
L'1~" ,
I I
. . . .
.£ -6~'. ............
- 100
-200
"0 o : 5 , . . . . . . . . . . . . . O. 0 Advanced [~] Argillic Alteration ] Montrnorillonite ]
Zone Zeolile Zone
] Epidole Zone
[ ~ Au-Ag Vein Figure 1.76. Schematic c o l u m n at section o f the Seigoshi A u - A g mine area, central J a p a n (Shikazono, 1985a).
temperature and pressure are 170-185°C at 2 - 5 kbar respectively. However, on the basis of experimental studies (Liou, 1971b; Maruyama et al., 1983) it is safe to expect that stilbite was formed under lower temperature conditions, probably less than ~150°C. Thus, it is considered that zone (3) was formed under temperatures less than ~ 150°C. In summarizing the fluid inclusion studies and stability of zeolite minerals, the most likely temperature range o f zone (1), (2) and (3) is estimated to be ~ 2 5 0 - 2 8 0 ° C , 150-230°C and < 150°C, respectively. Boiling of fluids for zones (2) and (3) suggests that the depth for the zeolite zone is probably less than 500 m from the surface and the epidote zone is more than 500 m.
Miocene-Pliocene Hydrothermal Ore Deposits
107
+2 +1 0
A
Broadlands . . t~ \(wel~ll) . _, . - " .~ • Mexixali (well5) Larderello \ \. . ~,; ~ ~, . . . " Nasu Chausudake ,, - , % Wairakei tlveragera- .%k\, ,\ ,,o . ' , ~.-,, • . . . . c ' Showashinzan ~ 0 * {average) :~:4[~ ~"
The G e y s e r s
~-~ 1 O 2
Showashinzan -'~,~, ;*" " '~ . ' " .~'./~huachapan .." - " El Tatio . -
3
Satsuma Iwojirna
...~':i~i~'~ : ::~
\B
-4
50
200
2;0
3;0
3;0
4;0
450
T(°C)
Figure 1.77. Range of hydrogen sulfide fugacity (fH2S) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica-rich zone) and some active geothermal systems. A = propylitic alteration; B = advanced argillic alteration. Data on active geothermal systems are taken from Ellis and Mahon (1977). The calculation of fH2s of Showashinzan for saturated water vapor pressure condition was based on the analytical data on volcanic gas by Matsuo (1961). The calculation of fH2S of Satsumalwo-Jima was based on the analytical data on volcanic gas by Matsuo et al. (1974) assuming PH20 = 0.5 kbar. ThermochemicaI data necessary for estimating fH2s for the propylitic and advanced argillic alterations are taken from Bird and Helgeson (1981) and Giggenbach (1981). The curves A-A', B-B' and C-C' represent the equilibrium of epidote(xp~ =0.30) - - K-mica(aK mica=0.9) - K-feldspar(a~ feidspar=0.95) -- pyrite - calcite - chlorite(aFco_0.5) where Xpis, aK-mica, aK-feldspar, and aFeo are mole fraction of pistacite component in epidote, activity of K-mica component in mica, activity of K-feldspar component in K-feldspar and activity of FeO component in chlorite, and hematite + liquid sulfur .~- pyrite + H2S, respectively (open circle = vapor-dominated system; solid circle = hot-water-dominated system; solid triangle = volcanic gas) (Shikazono, 1985a).
Ranges of $2, 02 and H2S fugacities were estimated as shown in Figs. 1.81, 1.82 and 1.77, respectively. $2 and 02 fugacities were estimated from the alteration mineral assemblage and homogenization temperature data. H2S fugacity was estimated from the chemical compositions of epidote and chlorite following the procedure by Giggenbach (1980, 1981) as shown in Fig. 1.77. CO2 fugacity can be inferred from the following chemical equilibrium relations (Bird and Helgeson, 1981): 3K-mica + 4calcite + 6quartz = 2clinozoisite + 4K-feldspar ÷ 4CO2 + 2H20
(1-29)
Based on the analytical data of K-mica, epidote and K-feldspar and using thermochemical data on these minerals (Helgeson and Kirkham, 1974; Helgeson et al., 1978; Bird and Helgeson, 1981), the fco2 range for the propylitic alteration was estimated (Fig. 1.78). The regional distribution of advanced argillic alteration in this district is shown in Fig. 1.73. The alteration zone of this type is distributed with a given trend, in general running from north to south. The distribution area of this type of alteration is more restricted than that of the propylitic alteration (Fig. 1.73). This type of alteration is well observed in the Funabara area, underground in the Seigoshi mine and the Ugusu silica mine (Figs. 1.73 and 1.79). The original rocks for this alteration are different in these
Chapter 1
108
T.Alfina I */Bagn/re @4
1
0 110
0
0 _1
-1
Broadlands /~" Nilan/d~Krafla (~zildere (~Ngawha ~ ~ 2 p h a s e Rotorua V O~.~J'~
~ ~'~'~'~~'~'r ~
"-~L~------''~2Krafl pha%e ~ \ "CerroPrieto Sa,,onSea
/ ~ ~ A '
,~,and
" Seigoshi
-2 -3
-4 I
[
I
I
150
200
250
300
T(°C)
Figure 1.78. Range of carbon dioxide fugacity (fco2) and temperature for the propylitic alteration (epidote zone) in the Seigoshi area and some active geothermal systems. Seigoshi = propylitic alteration of the Seigoshi district. Data on active geothermal systems are taken from Helgeson (1967, 1968), D'Amore and Panichi (1980), Giggenbach (1980, 1982), Bird and Norton (1981), and Arndrsson and Gunnlaugsson (1983). ThermochernicaI data necessary for estimating fco2 for the propylitic alteration are taken from Helgeson and Kirkham (1974), Helgeson et al. (1978), and Bird and Helgeson (1981). The curves A-B and A ' - B ' are equilibria for epidote(xp~=o._~o) - - K-miCa(aK,.~c~ o.9) - - K-feldspar(aKqeid~p~r-0.95) - - calcite assemblages at saturated water vapor pressure condition.
areas; Tanaba andesitic rocks (Funabara area), basic intrusive rocks, pyroclastic rocks of the Yugashima Group and Tanaba Andesite (Seigoshi area) and Koshimoda Andesite, and andesitic lava and pyroclastic rocks of the Yugashima Group (Ugusu area). The elevation level of this alteration is high, ~200-400 m higher than the top of the A u - A g vein. Forms of the alteration halo are also different in these three areas; lenticular and mushroom like (Funabara), vein (Seigoshi) and strata-bound and lenticular (Ugusu). Among these three areas, the most intense alteration has occurred in the Ugusu area. Thus the description on the alteration in this area is given below as an example of the advanced argillic alteration. The Ugusu silica mine, situated 4 km south of the Seigoshi mine (Fig. 1.73), produces 6 x 105 tons of silica ore annually, which contains more than 95 wt% SiO2. The geology and the lateral and vertical alteration zonings of this mine are shown in Figs. 1.79 and 1.80. Quartz is the most predominant phase in the central zone and the silica content of this zone is 95-99 wt%. Native sulfur is found in the pores of the highly silicified rocks together with small amounts of topaz. Alunite occurs abundantly, surrounding this zone, with quartz. Dickite, sericite and sericite/montmorillonite mixed-layer mineral occur in the more peripheral zone. Montmorillonite (tri-type, contrasting with di-type of the propylitic alteration) occurs in the most peripheral zone. Generally, the boundary between the alunite zone and surrounding clay zone is sharp, but the boundary between the silica zone and alunite zone is gradual. Pyrite occurs in the clay zone. Hematite
Miocene-Pliocene Hydrothermal Ore Deposits /i 400 ~ ',--/-'.~,; - 5OO
.
/
)\
,,/
109
...~'~'--.o~ Hakko
.
-
"x.J.
..
.....6~a.'~::
:ii:"'=ii "'':.
BB
::::: ::..:::
Silica-rich Zone Alunite-rich Zone
L
j
~
1kin
Clay-rich Zone
~2~ Unconformity
Figure 1.79. Geology and alteration zoning in the Ugusu silica mine (modified from Iwao, 1949, 1962). Toi F = Toi Formation; Koshimoda = Koshimoda andesite; Hakko = Hakko orebody; Shibayama = Shibayama orebody. Numbers indicate metres above sea level. Alteration zoning in the section of A'-B ~ is shown in Fig. 1.80 (Shikazono, 1985a)
Margin(A') Lower
Montmorillonite a Cristobalite Ser./Mont
.
Sericite Kaolin Alunite Topaz Native Sulfur Quartz Pyrite Hematite Rutile
.
.
.
.
.
Center(B') Upper
.
........ ___ ___
................ .................
Figure 1.80. Zonal sequence of the advanced argitlic alteration from the central to marginal zone in section of A'-B ~ in Fig. 1.79 and from upper horizon to lower horizon (Shikazono, 1985a)
o c c u r s as d i s s e m i n a t i o n a n d v e i n l e t s w i t h q u a r t z in t h e silicified a n d a l u n i t e zones. R u t i l e c o m m o n l y o c c u r s in the a l t e r a t i o n z o n e , e s p e c i a l l y in the c e n t r a l z o n e . A n a t a s e t e n d s to o c c u r in t h e m a r g i n a l p a r t o f the a l t e r a t i o n z o n e . S m a l l a m o u n t s o f o~-cristobalite a n d t r i d y m i t e o c c u r in the m a r g i n a l z o n e . T h e fluid i n c l u s i o n s c a n b e d i v i d e d i n t o t w o types: v a p o r - a n d l i q u i d - r i c h fluid i n c l u s i o n s . T h e filling d e g r e e o f fluid i n c l u s i o n s f r o m s o m e s a m p l e s f r o m the silicified a n d a l u n i t e z o n e s is v a r i a b l e a n d h o m o g e n i z a t i o n t e m p e r a t u r e s v a r y widely. T h i s i n d i c a t e s
Chapter 1
110
that the liquid-vapor separation occurred during the hydrothermal alteration process for the silica and alunite zones. Although the filling degree is variable, the homogenization temperature of the samples from the silica-rich zone is high, being in the range of 285-430°C; that from the alunite-rich zone in the range of 240-360°C and that from the alunite-clay-rich zone in the range of 220-280°C. This indicates that a steep temperature gradient existed in the hydrothermal alteration zones probably due to the mixing of high temperature volcanic gas and low temperature groundwater. Based on the hydrothermal alteration mineral assemblages and the fluid inclusion, the probable range of gas fugacities (fs2, fOR, fH2S) and temperature can be seen in Figs. 1.81 and 1.82: these estimated fugacities are quite different from those of the propylitic alteration. The characteristic features and differences of the two types of alteration are schematically summarized in Fig. 1.76, indicating the vertical changes in geology, alteration minerals and fluid inclusion characteristics in the Seigoshi-Ugusu district. The vertical alteration zoning from the shallower portion to the deeper portion in this district is summarized as follows: a massive part of the advanced argillic alteration zone; the zeolite zone of the propylitic alteration; and the epidote zone of the propylitic alteration. The advanced argillic alteration zone is developed near the unconformity boundary between relatively permeable pyroclastic rocks and the overlying relatively impermeable pyroclastic rocks. The upper portion of this district (zeolite and advanced argillic zones) is considered to be a two-phase (vapor and liquid) separation zone. It is worth noting that the mineralogical sequence in some active geothermal areas characterized by high chloride concentrations is similar to that found in the propylitic alteration of this district. For example, the alteration zoning in Wairakei, New Zealand, is similar to that of the Seigoshi district: in the shallower zone montmorillonite, illite-montmorillonite mixed-layer mineral, and laumontite occurs, and in the lower part
s~ c'd
co
f
s ~. "" j..-z.----~_ ~3,
-15 -20 -25
150
200
250
300
T (°C)
Figure 1.8i. Range of sulfur fugacity (.Ds2) and temperature for the propylitic alteration (epidote zone) and the advanced argillic alteration (silica- and alunite-rich zones (Shikazono, 1985a)). A = propytitic alteration; B = advanced argillic alteration. Thermochemical data on the fs2-temperature boundaries for the equilibria of: liquid sulfur ~- S2gas; hematite + pyrite ~- magnetite + S2gas; and pyrite ~ pyrrhotite + S2gas were taken from Heigeson (1969) and Rau et al. (1973a,b). S(1) = liquid sulfur; S2(v) = S 2 g a s ; ht = hematite, m t = magnetite; py = pyrite; po = pyrrhotite.
Miocene-Pliocene Hydrothermal Ore Deposits
tll
•
i.i i-
-30 -35 -40 04 O
"o
-45 -50 -55 -60
I
I
I
I
150
200
250
300
T (°C)
Figure 1.82. Range of oxygen fugacity (f%) and temperature for the propyiitic alteration (epidote zone) and the advanced argillic aIteration (silica- and alunite-rich zones (Shikazono, i985a)). A = propylitic alteration; B = advanced argilIic alteration. Thermochemical data on the fo2-temperature boundaries for the equilibria of: liquid sulfur + hematite ~ pyrite + H20 for saturated water vapor pressure condition; hematite ~ magnetite + O2; and pyrite + magnetite ~ pyrrhotite + 02, are taken from Helgeson (1969) and Rau et al. (1973a,b). Abbreviations used are the same as in Fig. 1.81. epidote and wairakite are found (Steiner, 1968). Both areas (Seigoshi and Wairakei) are characterized by c o m m o n occurrences of zeolites (laumontite and wairakite) and epidote and small amounts o f carbonates. The Larderello region (Italy) is also characterized by the occurrence of wairakite, laumontite and epidote (Cavaretta et al., 1982). Environmental conditions o f gas fugacities (fHzS, fco2) and temperature at the Wairakei and Larderello regions are similar to those of the Seigoshi district as shown in Figs. 1.77 and 1.78. The estimated depth o f the part affected by the zeolite and epidote alterations in the Seigoshi district is consistent with those of Wairakei and Larderello. The formation of epidote, K-feldspar, prehnite, wairakite and calcite in the geothermal area is considered to be due to the loss o f CO2 gas and rapid precipitation from the solution supersaturated with respect to quartz (Browne, 1978). The widespread occurrence o f these minerals in the Seigoshi district seems to be consistent with the above-mentioned consideration, namely that these minerals usually occur as veinlets rather than the replacements of original minerals and filling amygdule. In particular, many veinlets of epidote, prehnite and wairakite are found near the A u - A g - q u a r t z veins. There are three possible mechanisms for generating the strongly acid solution which caused the advanced argillic alteration: (1) alteration caused by the vapordominated system as inferred by White et al. (1971); (2) alteration caused by the oxidation o f HzS near the surface; and (3) alteration by volcanic gas a n d / o r hot water condensed from a volcanic gas. A m o n g them, (3) is the most attractive mechanism given the following evidence and considerations.
112
Chapter 1
(1) The temperature of the advanced argillic alteration estimated from the fluid inclusion studies and mineral assemblages, varies widely from ~220°C to ~420°C. The homogenization temperature for the central part of the alteration zone is from ~285°C to ~430°C. This temperature range is higher than the temperature of the vapor-dominated system defined by White et al. (1971) who showed that vapor-dominated systems such as Larderello (Italy), The Geysers and Mud Volcano (Yellowstone Park, Wyoming, USA) and Matsukawa (northeast Japan) are characterized by the temperature of around 240°C. (2) The advanced argillic alteration minerals are similar to those of the vapordominated system. But, in general, the minerals indicating relatively higher temperatures, such as pyrophyllite and diaspore, are lacking in the vapor-dominated system mentioned above. However, in Matsukawa which may be a vapor-dominated system at present, these high-temperature minerals like pyrophyllite, diaspore and andalusite have been reported (Sumi, 1968a). However, these minerals are considered to have been formed during the past and are not presently forming (Sumi, 1968b). (3) The minerals containing F and C1 such as topaz and zunyite are common in the advanced argillic alteration. The HF activity of the solution was estimated to be ~0.01 based on the F content of topaz from the Ugusu mine (Shibue and Iiyama, 1984). This high F concentration is observed in Crater Lake (Mt. Ruapehu, New Zealand) (Giggenbach, 1974) and is interpreted to be due to the injection of acid fumarolic gases to the lake (Giggenbach, 1974). Gases collected from a fumarole at Showashinzan and Satsumaiwo-Jima also have a high H F / H 2 0 volume ratio which is in the range of (2-5)× 10 -4 (Matsuo et al., 1974; Mizutani and Sugiura, 1982). (4) If alunite, K-mica and kaolinite (which are common minerals in the advanced argillic alteration) are in equilibrium, the concentration of H2SO4 can be estimated based on the experimental work by Hemley et al. (1969); the concentration of H2SO4 at 200°C and 300°C is 0.002 and 0.012 M, respectively. This may suggest that it is difficult to form such a high concentration of sulfate ion only by oxidation of H2S. (5) The fugacity of H2S (fH2s) for the advanced argillic alteration is plotted in Fig. 1.77 together with those for some active geothermal systems and volcanic gas. The fH2S for the advanced argillic alteration was estimated on the basis of hematite-pyriteliquid sulfur equilibrium. Estimated fH2S for the advanced argillic alteration is lower than that for the vapor-dominated system (Larderello and The Geysers) but is similar to that of the volcanic gas collected from one of the fumaroles of the Showashinzan volcano in Japan. Only a few areas subjected to solfataric alteration have been well studied. Satsumaiwo-Jima is probably the only area where alteration caused by volcanic gas has been studied in detail. The zonal sequence from the centre to the margin and from a topographically higher level to a lower level is from a silica-rich zone through an alunite-rich zone to a clay (montmorillonite)-rich zone (Yoshida et al., 1976). This pattern is similar to that of the advanced argillic alteration in the Seigoshi-Ugusu district, although oe-cristobalite is abundant and quartz is poor in the Satsumaiwo-Jima. But in the Seigoshi oe-cristobalite is poor and quartz is abundant. In the Satsumaiwo-Jima pyrophyllite, diaspore, zunyite and topaz are not found, although these minerals occur in the Seigoshi-Ugusu district. The occurrence of alunite in the Satsumaiwo-Jima indicates that the formation of alunite is largely controlled by the existence of groundwater. The
Miocene-Pliocene Hydrothermal Ore Deposits
l 13
alunite zone in this island is strata-bound in form. In the Ugusu mine the occurrence of alunite is very similar. It is interesting to note that the advanced argillic alteration tends to occur near the unconformity boundary (Fig. 1.79). The similarity in the mode of occurrence of alunite from the Satsumaiwo-Jima and from the Ugusu mine, the existence of the unconformity and the steep temperature gradient from central to marginal part in the Ugusu mine may suggest that the mixing of groundwater with volcanic gas and/or hot water condensed from the volcanic gas is the most important mechanism for the formation of advanced argillic alteration also in the Satsumaiwo-Jima area. The coexistence of these two types of hydrothermal alteration and associated high sulfidation-type A u - A g deposits (acid alteration) and low sulfidation-type Au Ag deposits (neutral alteration) is observed not only in the Seigoshi-Ugusu district, but also in the other epithermal A u - A g vein-type mine districts in Neogene volcanic region in Japan: Izu Peninsula, Takadama, Kushikino, Hishikari, Akeshi-Kasuga, Yatani, and Harukiyama A u - A g mine (Hokkaido) (Yamada, 1995) districts. In these districts both the epithermal vein-type deposits and highly silicified and clay-rich rocks are found in close proximity in time and space. The characteristic differences between the two types of coexisting associated alterations in each area are very similar to those in the Seigoshi-Ugusu district. For instance, advanced argillic alteration zone lies at topographically higher levels and in younger formations than the epithermal vein-type deposits and associated propylitic alteration. However, some differences exist in a different area. For instance, in the Kasuga-Akeshi area, fine Au particles occur in the highly silicified rocks. Surrounding the highly silicified rocks, the following zonal sequence is observed from the centre to the margin (Tokunaga, 1955; Saito and Sato, 1978): an alunite zone, kaolinite zone and a montmorillonite zone. This sequence is similar to that observed in the Ugusu deposits, but Au has not been detected in the silicified part of the Ugusu silica deposits. Topaz and zunyite have not yet been reported from the Kasuga silicified rock, while they are found in that of the Ugusu. Homogenization temperatures for the silicifled rocks in the Kasuga, Akeshi, Iwato and Kushikino areas are generally lower than those for the Ugusu deposits (Takenouchi, 1981). Boiling phenomena have been observed in the fluid inclusions from the silicified rocks in these areas (Takenouchi, 1981). Yamada (1995) described acid and potassium alterations and high sulfidation and low sulfidation mineralizations in the Harukiyama district (Hokkaido). He indicated based on K - A r data on the alterations that the high sulfidation-type have been formed first by the residual magmatic fluid subsequent to the intrusion of quartz porphyry magma and then the low sulfidation-type could have been formed by the circulation of meteoric water generated by the heated emanation from hot magma solidified but still similar temporal relationship is observed in the other districts (Osorezan: Aoki, personal communication, 1990; Yatani: Shikazono, unpublished; Seigoshi-Ugusu: Shikazono, unpublished).
1.4.2.6. Chemical composition of alteration minerals Although a wide range of alteration minerals has been recognized in epithermal systems considered here, few of their chemical compositions have been determined. Trioctahedral chlorite occurs commonly in geothermal and hydrothermal areas, whereas the occurrence of dioctahedral chlorite is very limited. For instance, donbasite
114
Chapter I
does not occur in geothermal and hydrothermal areas, but sudoite (Al-chlorite) commonly does not occur in Kuroko mine area (e.g., Tsuzuki and Honda, 1977). Dioctahedral chlorite has not been reported from the Neogene C u - P b - Z n vein-type deposits in Japan; instead, trioctahedral chlorite is common (Shirozu et al., 1975). In Kuroko mine area, the common chlorite is trioctahedral Mg chlorite. The structural formula for trioctahedral chlorite is represented by (Mg6-x-yF@ + Alx)(Alx Si4-x)O10(OH)8. It is convenient to plot a diagram of Fe2+/(Fe 2+ + Mg) (in atomic fraction) against ~VA1/(Si + IVA1)to display compositional variations in chlorite (Fig. 1.83) (e.g., Hey, 1954; Foster, 1962, Nagasawa et al., 1976). Although a large body of analytical data on hydrothermal chlorite from geothermal and hydrothermal areas is available, there are few data on the chemical composition of original fresh volcanic rocks. The relationship between the ratio MgO/FeO in chlorite and that in the fresh host-rocks indicates that the value of MgO/FeO of chlorite generally satisfies a line of 1 : 1 slope. The correlation between MgO/FeO in the host rock and that in the chlorite implies that the MgO/FeO value of chlorite from propytitically altered rocks associated with the mine areas and from altered rocks in terrestrial and submarine geothermal areas is largely affected by MgO/FeO ratio of original fresh rocks. However, most of the data from Kuroko and Neogene C u - P b - Z n vein-type deposits deviate significantly from this line. Chlorite compositions from areas (Toyoha Pb-Zn vein, Kuroko deposits) deviate significantly from a line of 1 : 1 slope. This deviation implies that the Fe2+/Mg value of chlorite from these areas is controlled not only by the FeO/MgO value of the fresh host rocks, but also by factors such as the ratio of Fe 2+ to Mg 2+ in the fluid phase. As noted already, several investigators have shown isotopically that seawater played an important role in the formation of Kuroko deposits (e.g., Sakai et al., 1970; Kajiwara, 1971; Hattori and Sakai, 1979; Farrell and Holland, 1983). Mg-rich chlorite occurs in gypsum-anhydrite bodies in many Kuroko deposits. Farrell and Holland (1983), Shikazono et al. (1983), and Kusakabe and Chiba (1983) suggested the involvement of large amounts of seawater or seawater-dominated hydrothermal solution in the formation of the gypsum-anhydrite bodies. It is reasonable to assume from the large number of experimental studies on rock-seawater interaction at elevated temperatures that the seawater-dominated fluid phase, which interacted with volcanic rocks at a high water/rock ratio, contained appreciable amounts of Mg but very small amounts of Fe (Seyfried and Mottl, 1982). Therefore, it is likely that Mg-rich chlorite precipitated from a solution with a high proportion of Mg 2+ to Fe 2+. In contrast to the Kuroko hydrothermal system, there is no evidence for the involvement of seawater in the hydrothermal system associated with the C u - P b - Z n vein mineralization. Results of hydrogen and oxygen isotopic studies indicate that large amounts of meteoric water were incorporated into the ore fluids responsible for these vein-type deposits (Hattori and Sakai, 1979). If F e - M g chlorite is assumed to be in equilibrium with a fluid phase and pyrite, the ratio of Fe 2+ to Mg 2+ in the fluids may be related to factors such as pH, f Q , temperature and total dissolved sulfur concentration (ES). This relationship can be derived from the following chemical reactions: MgsA1Si3A1Olo(OH)8 + 5 Fe 2+ = FesA1Si3A1Olo(OH)8 + 5 Mg 2+
(1-30)
Miocene-Pliocene Hydrothermal Ore Deposits
115
1.0
oI 03 01
04
0.8
05
._.0.6
o)
o 6
oO6
4LL
o6
v
o6
u_ 0.4
°6o6
o6
0.2
9 ~ 910_10 • •8 o 8 8 o8
09
0.0
018
1:0
1:2 1:4 AI in 4(AI,Si)
o7o 7
1:6
1:8
2~0
Figure 1.83. Variation of Fe2+/(Fe 2+ -}- Mg) and tetrahedral A1 of chIorite from hydrothermal ore deposits: Japanese Neogene Cu-Pb-Zn vein-type (open circle) and Kuroko deposits (solid circle). Localities: 1: Ashio (Nakamura, 1960, 1963); 2: Yatani (Hattori, 1974); 3: Toyoha (Shikazono 1974a, Sawai, 1984); 4: Kishu (Shirozu, 1958); 5: Sayama (Shirozu, 1958); 6: Mikawa (Nagasawa, 1961); 7: Furutobe (Shirozu et al., 1975); 8: Hanaoka (Hayashi 1961, Hayashi and Oinuma, 1965; Tsuzuki and Honda, 1977; Shirozu et aI., 1975); 9: Wanibuchi (Sakamoto and Sudo 1956, Iwao and Minato 1959, Katsumoto and Shirozu, 1973); 10: western Bergslagen (Baker et al., 1983) (Shikazono and Kawahata, 1987).
FeS2 + 2 H + + H 2 0 = F e 2+ + 2H2S + 1 / 2 0 2
(1-31)
w h e r e MgsA1Si3A1OIo(OH)8 and Fe5A1Si3A1Olo(OH)8 represent M g - c h l o r i t e and Fechlorite, respectively. Generally, activities o f liquid H 2 0 and FeS2 do not deviate f r o m unity; thus these values are a s s u m e d to be unity.
116
Chapter 1 From equations (1-30) and (1-31), we obtain, for the H2S dominant region,
log(aFe-chl/aMg-chl) = log K1-30 + 5 log Kl-31 -- 5 log aMg2+ -- 10 log ES - 10 log gHzS -- (5/2)1og fo2 -- 10pH
(1-32)
where g is the activity coefficient. For the SO 2- dominant region,
1og(aFe-chl/aMg-chl) = log K ~_30+ 5 log K 1-31 - 5 log aMg2+ 10 pH -- 10 log ES - 10 log YSO2- + 35/2 log fo2 + 10 log KI-34
(1-33)
where KI_34 is the equilibrium constant for the following reaction, H2S+202=SO ] +2H +
(1-34)
Equations (1-32) and (1-33) imply that the aFe--chl/aMg-chl of chlorite depends on foz, aMgz+, ~S, pH, temperature and ionic strength. From equations (1-32) and (1-33), it can be shown that the aFe-chl[aMg-chlof chlorite in equilibrium with pyrite decreases with increasing fo2 at constant temperature, pH, aMg2+, ionic strength, and ES in the region where H2S is dominant, whereas it increases with increasing f Q in the SO ] - dominant region. Previous studies on the estimates of f Q , fs2, pH, ZS and temperature for Kuroko ore deposition have been reported, for instance, by Kajiwara (1971), Shikazono (1976) and Ohmoto et al. (1983), who showed that the fo2 of the Kuroko ore fluids lies in the region close to the SO42-/H2S boundary (in this expression, SO ] and H2S represent the concentration of the total dissolved oxidized sulfur species and total dissolved reduced sulfur species, respectively). These estimates seem consistent with the tow Fe2+/Mg value of chlorite from Kuroko deposits. Chlorite from the Toyoha Pb-Zn vein type deposits is associated with sphalerite, pyrite and, rarely, pyrrhotite (Shikazono, 1974a). The iron content of the sphalerite associated with chlorite is 1.2-2.9 wt%. The temperature of formation of chlorite in the Toyoha deposits is estimated to be 200-250°C from fluid inclusion data (Shikazono, 1974a, 1975; Yajima and Ohta, 1979), which indicates that the chlorite was formed in a relatively reducing environment in which reduced sulfur species predominate. This estimate seems consistent with the composition of chlorite from the Toyoha deposits and from the other Neogene C u - P b - Z n vein-type deposits in Japan, which contains up to 40 wt% FeO (Nakamura, 1960, 1963; Shirozu, 1969; Shikazono, 1974a; Hattori, 1974). Shikazono (1974a,c, 1977a, 1978b) and Hattori (1975) showed that these deposits formed in the environments where species of aqueous reduced sulfur predominated. This estimated range of oxidation state and ratio of concentration of aqueous reduced sulfur species to oxidized sulfur species appears to be in agreement with the chemistry of chlorite from these deposits. However, the Fe2+/Mg of chlorite also depends on the other factors, such as IES, pH and aMg2+. From equations (1-32) and (1-33), it is obvious that increasing IgS and aMg2+ also causes a lower Fe2+/Mg in chlorite. The Fe 2+ /Mg and Fe-3+ /Fe 2+ values of chlorite from Kuroko deposits and Neogene C u - P b - Z n vein-type deposits differ greatly (Fig. 1.83). Chlorite from Kuroko deposits contains lower Fe2+/Mg and higher Fe3+/Fe 2+ values than this from the Neogene vein-type deposits in Japan. The most likely explanation for these differences is that these two types of deposit formed at different states of oxidation, although other
Miocene-Pliocene Hydrothermal Ore Deposits
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variables such as pH, temperature, aMg2+ and NS are also possible important factors, as discussed above. Generally, the Fe z + / M g value of chlorite from a given geothermal or hydrothermal area is variable. The Fe2+/Mg value of chlorite in the Toyoha mine district varies widely. Sawai (1984) has shown that the iron content of chlorite away from the P b - Z n veins. Iron content of chlorite in the Toyoha veins is very high (40 wt% FeO: Shikazono, 1974a; Sawai, 1984). Variations in iron content of chlorite from the host rocks toward the C u Pb-Zn veins have also been studied for other Neogene vein-type deposits in Japan (e.g., Ohe, Ashio deposits: Hayashi, 1979). Variations in iron content of chlorite from Japanese epithermal mine districts suggest that the iron content of chlorite in the discharge zone of a hydrothermal system is higher than that in the recharge zone. Mottl (1983) has also suggested that the FeZ+/Mg value of chlorite in the discharge zone of submarine geothermal systems is higher than that in the recharge zone. The Fe2+/Mg of chlorite from different parts of a geothermal system is higher than that in the recharge zone. The FeZ+/Mg of chlorite from different parts of a geothermal system (recharge zones versus discharge zones) is probably not constant. In order to evaluate the effect of the fluid movement at discharge versus recharge zones, the temperature dependence of Fe2+/Mg 2+ in fluids in equilibrium with the isochemically recrystallized crystal rocks will be considered below. By taking a value for Fe2+/Mg of chlorite that is equal to that of the average andesitic and basaltic rocks (0.6), and assuming that chlorite is an ideal solid solution of 14 A Fe-chlorite and 14 A Mg-chlorite, the dependence of a aFez+/aMg2+in fluids on temperature was calculated by using thermochemical data for chlorite from Walshe and Solomon (1981). Figure 1.84 shows that aFez+/aMg2+of fluids increases with increasing
-2
+-3 04
~-4 o
200
250
3()0 350 Temp.(°C)
Figure 1.84. Variation of aFe2+/aMg2+of hydrothermal solution in equilibrium with chlorite having constant Fe2+/Mg (=0.6), as a function of temperature. The significance of points A and B is discussed in the text (Shikazono and Kawahata, 1987).
118
Chapter I
of temperature. Therefore, if fluids initially in equilibrium with chlorite having the same FeO/MgO as that of average andesitic and basaltic rock at elevated temperature (for example, at point A in Fig. 1.84) ascend rapidly without interaction with the surrounding rocks, chlorite precipitating from fluids at lower temperature (for example, at point B in Fig. 1.84) could contain appreciable amounts of Fe 2+ compared with Mg. This mechanism could lead to the formation of chlorite having an unusually high content of iron. It is also noteworthy that the vein chlorite in the altered basalt from the Costa Rica Rift contains higher concentrations of iron than the chlorite that replaces mafic minerals in the rock (Kawahata, 1984). In this case it is likely that the flow rate of ascending fluids from which vein chlorite precipitated is high compared with the rate of reaction between the ascending fluids and surrounding rocks. The Fe2+/Mg value of chlorite precipitating from ascending fluids depends on the extent of deviation from equilibrium between fluids and surrounding rocks. As discussed in detail by Giggenbach (1984), a number of processes such as adiabatic and conductive cooling of fluids and mixing of fluids can cause this deviation. The above considerations suggest that chlorite occurring in the discharge zones of hydrothermal systems would contain a higher concentration of iron than that occurring in recharge zones. If a magnesium-rich solution such as seawater or a seawater-dominated hydrothermal solution was involved in the hydrothermal systems, it could be expected that chlorite, even that occurring in discharge zones, would contain a high content of Mg. It is widely accepted that ascending hydrothermal solutions mixes with cold seawater at the time of ore formation at Kuroko (e.g., Hattori and Sakai, 1979; Shikazono et al., 1983). The Mg concentration of this hydrothermal solution increased with higher degrees of mixing. Involvement of large amounts of seawater at the site of ore deposition could be one of the reasons why chlorite from Kuroko deposits contains high amounts of Mg. In addition to a large concentration of Mg in these fluids, relatively high fo2, NS and SO42-/H2S, and low pH, could lead to the large Mg content of chlorite in Kuroko deposits, as discussed above. Consequently, the composition of chlorite in the discharge zone depends largely on the chemical nature of fluids (factors such as Fe2+/Mg 2+, SO42-/H2S, pH, aMg2+) and temperature. Movement of fluids may also be an important cause for the variability in the ratio of Fe 2+ to Mg in hydrothermal chlorite. Wide compositional variations in chlorite from the hydrothermal ore deposits in Japan, including Kuroko and Neogene C u - P b - Z n vein-type deposits, are considered to reflect the variable chemical nature of ascending ore fluids and fluids that mix with ascending ore fluids at discharge zone. The variations in Fe and Mg contents of the 14 ,~ Fe-chlorite-14 ,& Mg-chlorite solid solution are considered here. However, structural formulae for chlorite are not as simple as those considered here. As mentioned by Walshe and Solomon (1981), Stoesell (1984), Cathelineau and Nieva (1985) and Walshe (1986), chlorite solid solution may be represented by six components, and accurate thermochemical data on each end-member component at the hydrothermal conditions of concern are necessary to provide a far more rigorous calculation of the equilibrium between chlorite and hydrothermal solution. However, the above argument demonstrates that the composition of chlorite is a highly useful indicator of physicochemical conditions of hydrothermal solution and extent of water-rock interaction.
Miocene-Pliocene Hydrothermal Ore Deposits
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Epidote is one of the most common alteration minerals occurring in geothermal and mine areas. The factors controlling the chemical composition of epidote from geothermal areas have been examined from thermochemical points of view. Bird and Helgeson (1981) discussed the effects of temperature, CO2 and 02 fugacities, and activity ratios (aFe3+)/(aH+) 3, (aca2+)/(au+) 2, and (aai3+)/(aH+) 3 in geothermal waters on the variations in iron and aluminum contents of epidote. Giggenbach (1981) calculated the effects of partial pressure of CO2 gas and temperature on iron content of epidote in equilibrium with alteration minerals such as kaolinite, K-mica, K-feldspar, albite, paragonite and calcite in geothermal systems. D'Amore and Gianelli (1983) showed the dependence of iron content of epidote on 02 fugacity for the epidote-K-feldspar-albitetremolite-chlorite equilibrium assemblage. Wolery (1978) and Reed (1982, 1983) have indicated based on a computer calculation of the change in chemistry of aqueous solution and mineralogy during seawater-rock interactions that epidote is formed under the low water/rock ratio less than ca. 50 by mass. Humphris and Thompson (1978), Stakes and O'Nell (1982) and Mottl (1983) have also suggested on the basis of their chemical and oxygen isotopic data of the altered ridge basalts that epidote is formed by seawater-basalt interaction at elevated temperatures (ca. 200-350°C) under the rock-dominated conditions. If epidote can be formed preferentially under such low water/rock ratio, the composition of epidote should be influenced by compositions of the original fresh rocks. Shikazono (1984) summarized analytical data of the epidote from geothermal areas to consider the relationship between the composition of epidote and that of the original fresh rocks and to inspect the other factors controlling the compositional variations in epidote. The discussion on the epidote composition by Shikazono (1984) is described below. Chemical compositions of epidote and original rocks in several geothermal and mine areas including Seigoshi (Japan, epithermal Au-Ag vein-type mine area), Yugashima (Japan, ancient geothermal area), Furutobe (Japan, Kuroko mine area), Ohtake (Japan, active geothermal area), Mid-Atlantic ridge, Costa Rica rift, Mitsuishi (Japan, pyrophyllite deposits), Shimokawa (Japan, ancient ophiolite area associated with Besshitype deposits), Reydarfjordur (Iceland, ancient geothermal area), Kushikino (Japan, epithermal Au-Ag vein-type mine area), Hachimantai (Japan, active geothermal area), Broadlands (New Zealand, active geothermal area), Wairakei (New Zealand, active geothermal area), Larderello (Italy, active geothermal area), and Sarmiento ophiolite complex (Chile, ancient ophiolite suite). In individual districts, these rocks are generally subjected to intense hydrothermal alteration. The original rock composition was estimated from analytical data of fresh rocks in the same area. Figure 1.85 shows that the Fe203 content of epidote positively correlates to Fe203 content of original rocks, although in general Fe203 content of epidote from each geothermal area has a variation over several wt%. Therefore, it is inferred that Fe203 content of original rocks has a large influence on iron content of epidote. Original rocks in Fig. 1.85 are mainly andesitic and basaltic rocks of island arcs and ocean ridge. In Fig. 1.85 iron contents of epidote from two different geologic environments, island arc and oceanic ridge or ophiolite, are summarized. It can be seen in Fig. 1.85 that the iron content of epidote from ridge basalt and ophiolite is generally lower than
Chapter 1
120
S
M. q
~13 O
2'
u..
1-
d"
sht
F,L=--~ Y
M, o ,cT J
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 Fe20s Content of Epidote (weight
%)
Figure 1.85. Relation between Fe203 content of epidote and that of original fresh rocks. S: Seigoshi, Y: Yugashima, N: Noya, F: Furotobe, O: Ohtake, M: Mid-Atlantic ridge, C: Costa Rica rift, Mi: Mitsuishi, Sh: Shimokawa (Shikazono, 1984)
that from the volcanic rocks of island arcs. One of the most likely explanations for this difference is the difference in Fe203 content of original rocks; Fe203 contents of oceanic basalts and ophiolite are generally lower than those of island arc volcanic rocks. The positive correlation shown in Fig. 1.85 is consistent with the considerations on seawater-volcanic rock interactions under the hydrothermal conditions by Wolery (1978), Humphris and Thompson (1978), Reed (1982, 1983), Stakes and O'Neil (1982) and Mottl (1983). They suggested that epidote is formed under a low water/rock ratio less than ca. 50 by mass. Mineralogical changes resulted from seawater-volcanic rocks interactions have been elucidated by these recent theoretical studies. The epidotes except those from the Furutobe Kuroko mine area, Mid-Atlantic ridge, Costa Rica rift, and Shimokawa are suspected to have been formed in subaerial geothermal environments. Thus it is uncertain in these cases that the epidote tends to form at a low water/rock ratio. However, it seems likely that the Fe/A1 ratio of epidote formed in subaerial geothermal areas is also reflected by that of original rocks because the concentrations of Fe 3+ and A13+ in geothermal waters interacted with the rocks under the hydrothermal condition are generally very low except for strongly acid and high chloride solutions. Iron content of epidote depends on kind of iron minerals coexisting with epidote. The iron content of epidote coexisting with hematite is relatively high. For example, the epidote coexisting with hematite from the Yugashima, Furutobe and Reydarfjordur has higher iron content than that from the other districts. This relation holds even in geothermal systems (Salton Sea: Keith et al., 1968; Furutobe: Shikazono et al., 1995). On the other hand, the epidote coexisting with pyrite and pyrrhotite contains generally smaller amounts of iron as observed in the Larderello (Cavaretta et al., 1982), Shimokawa, MidAtlantic ridge, and Costa Rica rift. This difference in the iron content of epidote coexisting with different opaque minerals could be explained by the difference in oxygen fugacity. The iron content of epidote coexisting with prehnite from the Seigoshi district is smaller than that without prehnite. It has been clarified that prehnite is stable at the low
Miocene-Pliocene Hydrothermal Ore Deposits
121
CO2 fugacity and that the iron content of epidote in equilibrium with prehnite is lower than that in equilibrium with other minerals such as K-feldspar, K-mica and calcite under such low CO2 fugacity conditions (e.g., Cavaretta et al., 1982). Therefore, it seems clear that iron content of epidote is affected also by CO2 fugacity. Although Fe203 content of epidote from a given geothermal area has roughly a positive correlation to that of original rocks, the variation in Fe203 content of epidote from a given geothermal area ranges over several weight percents. This variation is too large to be explained only in terms of the difference in Fe203 content of original rocks because the composition of epidote varies sometimes so widely even in a single grain. Compositional zoning in epidote grains from subaerial and island arc geothermal systems is commonly known; usually, iron content of epidote increases from core to margin in a crystal. This tendency is observed in the Seigoshi, Furutobe, Yugashima, and Larderello areas, even though a decrease in the iron content from core to margin is rarely recognized in a single grain from the Larderello (Cavaretta et al., 1982). Complicated compositional zonings of epidote from the Salton Sea geothermal area are also found (McDowell and McCurry, 1978). In contrast, epidote from midoceanic ridge basalts (Melson and Van Angel, 1966; Humphris and Thompson, 1978; Kawahata, 1984) shows no wide compositional variation in a single grain. Epidote from a typical midoceanic ridge basalt has a small variation range of iron (Humphris and Thompson, 1978). The variation range in epidote from the Costa Rica Rift studied by Kawahata (1984) is wider than that observed by Humphris and Thompson (1978). Kawahata (1984) considered that the alteration minerals including epidote in the DSDP Hole 504B in the Costa Rica rift were formed by the interaction of midoceanic ridge basalt and ascending hot fluids. The basalt in the Shimokawa area is probably an abyssal tholeiite (Kohsaka, 1975). In this area, strata-bound cupriferous pyritic deposits are found and general features of these deposits resemble those of the ore deposits of Cyprus type and the ore deposits found in East Pacific Rise 21°N. Therefore, the wide variation range in epidote from this area may be explained by the effects of the chemical compositions and gaseous fugacities of ascending hot fluids responsible for the ore formation together with the chemical compositions of original basalt. The compositional zoning of an epidote grain and wide compositional variations in epidote grain from each geothermal area are difficult to explain because iron content of epidote depends on many factors other than the chemical composition of original rocks. However, the differences in the variation range and compositional zoning of epidote from the subaerial and oceanic ridge geothermal systems suggest that the change of physicochemical variables such as oxygen fugacity and temperature during the hydrothermal activities in both geothermal systems is different. Few data on the chemical compositions of feldspars (albite, K-feldspar) are available. Fujii (1976) indicated that K-feldspar and albite in the propylite of west Izu Peninsula, middle Honshu are of nearly end member composition. Nagayama (1992) showed that K-feldspars in the Hishikari A u - A g vein and in the host andesitic rock have different composition; N a / K ratio of K-feldspars from the vein is lower than that from the host rocks.
Chapter 1
122
1.4.2.7. Causes for hydrothermal alteration Hydrothermal alteration is reflected by the changes in many variables (temperature, water/rock ratio, extent of water-rock interaction (reaction progress), reaction rate, flow rate of fluids etc.) (Fujimoto, 1987). Theoretical and experimental works on hydrothermal alteration were reviewed by Meyer and Hemley (1967), and Rose and Burt (1979). In the last two decades, great progress has been made in the field of hydrothemal alteration studies, mainly from computation works on water-rock interactions at elevated temperatures (e.g., Wolery, 1978; Reed, 1983, 1997; Takeno, 1989). These studies revealed the relationship between the changes in chemical composition of hydrothermal solution and the relative abundance of minerals in the rocks. There are different approaches to the study of hydrothermal alteration. For instance, Shikazono (1978a) showed the relationship between chemical composition of hydrothermal solution in equilibrium with the alteration minerals and C1- concentration in hydrothermal solution. Giggenbach (1984) calculated the effect of temperature on the chemical composition of fluids buffered by alteration minerals. The causes for the hydrothermal alteration considered below are mainly based on the works by Shikazono (1978a) and Giggenbach (1984). The effect of the extent of water-rock interaction is not taken into account. Figure 1.86 illustrates the variations in the chemical composition of chloride-rich hydrothermal solution in equilibrium with common alteration minerals with temperature. Figure 1.86 demonstrates that (1) the chemical compositions of hydrothermal solution
B
Na
H
E~ 0
Ca 2.
----4. i
-1
A
~
,-
I
-2
-3
~
150
200
r
250
t
300
Temp.(°C ) Figure 1.86. Variation in chemical compositions (in molaI unit) of hydrothermat solution with temperature. Thermochemical data used for the calculations are from Helgeson (1969). Calculation method is given in Shikazono (1978a). Chloride concentration in hydrothermal solution is assumed to be i moI/kg H20. A-B: Na + concentration in solution in equilibrium with low albite and adularia, C-D: K + concentration in solution in equilibrium with low albite and adularia, E-F: H4SiO4 concentration in equilibrium with quartz, G-H: Ca 2+ concentration in equilibrium with albite and anorthite (Shikazono, 1978a, 1988b).
Miocene-Pliocene Hydrothermal Ore Deposits
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depend on alteration minerals, temperature and C1- concentration; (2) K + and H4SiO4 concentrations increase with an increase in temperature, while Na + concentration does not largely depend on temperature. Calculations were made, assuming that the dominant anion is C1-. From these calculations, it is considered that the potassic alteration occurs when hydrothermal solution initially in equilibrium with propylitic alteration minerals ascends rapidly and interacts with country rocks at lower temperature ( I - J - K in Fig. 1.86). In this case, addition of K + to the rock takes place. K-bearing minerals such as sericite and K-feldspar precipitate from the fluid accompanied by the destruction of plagioclase in the country rocks and liberation of Ca, Sr, and Na to the fluids. Dissolutions of anorthite and albite components in a plagioclase occur by this mechanism. Thus, it is likely that Ca, Sr and Na are extracted from the rocks. It is expected that SiO2 content of the country rocks increases with progressive alteration because solubility of SiO2 decreases with a decrease in temperature ( O - P - Q in Fig. 1.86) (e.g., Holland and Malinin, 1979). If fluids initially in equilibrium with quartz ascend rapidly, some metastable minerals (amorphous silica, cristobalite, wairakite) may precipitate because of supersaturation with respect to SiO2 (e.g., Wolery, 1978; Bird and Norton, 1981). Important processes for the supersaturation and deviation from the equilibrium between fluids and rocks are adiabatic boiling, mixing of fluids and conductive cooling of fluids (Giggenbach, 1984). Formation of albite which is characteristic mineral of propylitic alteration occurs by heating of rocks and descending fluids at recharge zone in the hydrothermal system (Giggenbach, 1984; Takeno, 1989). Thus, it is considered that the propylitic alteration takes place at recharge zone in the hydrothermal system, while potassic alteration at discharge zone. Besides the effect of temperature, boiling and degassing play an important role for potassic alteration. Degassing of CO2 is accompanied by an increase in pH, causing the depositions of adularia and calcite (Browne, 1978). Probably, formations of some zeolite minerals (wairakite, yugawaralite, taumontite) are also caused by degassing and boiling of fluids and increase in pH (Shikazono, 1985a). There are two important chemical reactions which cause intermediate argillic and advanced argillic alterations H2S + 202 = 2H + + SO]
(1-35)
4SO2 + 4 H 2 0 = H2S + 3 H2SO4
(1-36)
Oxidation of H2S (reaction (1-35)) occurs under the near-surface environment. Oxygen may be supplied from oxygenated groundwater. These oxidation reactions liberate H + ion, leading to a decrease in pH. Under low pH conditions intermediate argillic alteration minerals (e.g., kaolinite, sericite) are stable. When temperatures of volcanic gases containing SO2 decrease, the reaction (1-35) proceeds to the right hand side. This reaction causes a considerable decrease in pH due to the formation of sulfuric acid. Advanced argillic alteration is formed by the interaction of volcanic gas with groundwater. The above interpretation on the alteration zoning is mainly based on thermodynamics. However, it is necessary to consider the influence of kinetics and fluids flow on
124
Chapter 1
the hydrothermal alteration processes to interpret the precipitations of metastable phases such as cristobalite. The coupled precipitation kinetics-fluid flow model was applied to the distribution of SiO2 content and K20 content of the hydrothermally altered andesite in the Hishikari A u - A g mine area, south Kyushu, Japan by Shikazono et al. (2002). This will be described in section 1.4.6. 1.4.3. Geochemical characteristics
Numerous geochemical data (fluid inclusions, stable isotopes, minor elements) on the epithermal vein-type deposits in Japan are available and these data can be used to constrain geochemical environment of ore deposition (gas fugacity, temperature, chemical compositions of ore fluids, etc.) and origin of ore deposits.
1.4.3.1. Fluid inclusions Substantial amounts of homogenization temperature data on the Neogene vein-type deposits in Japan are available (e.g., Enjoji and Takenouchi, 1976; Shikazono, 1985b) and they are summarized in Fig. 1.87. Homogenization temperatures vary widely within a given deposit type and even within a single deposit. However, the range of homogenization temperatures differs according to the type of deposit: 190°C 145 250oc for Au-Ag-rich deposits, 200°C to 250°C for Pb-Zn-Mn-rich deposits, and 200 ° to 350°C for Cu-Pb-Zn-rich deposits (Fig. 1.87) (Shikazono, 1985b). Homogenization temperatures are not same as the formation temperature. Therefore, we need pressure correction to estimate the formation temperatures from homogenization temperatures. However, the homogenization temperatures are in good agreement with electrum-sphalerite temperatures (Shikazono, 1985d) (Figs. 1.88 and 1.89). Therefore, pressure corrections to homogenization temperatures of fluid inclusions necessary to obtain formation temperatures are relatively small (less than 30°C). Salinities of inclusion fluids from epithermal vein-type deposits clearly indicate that the salinities of inclusion fluids from these types of deposits are distinctly different, that is, 20-2 NaC1 equivalent wt% (base-metal vein-type deposits) and 0-3 wt% (Au-Ag vein-type deposits) (Shikazono, 1985b) (Table 1.13). Salinities of inclusion fluids from Kuroko deposits (0.5-5 wt% NaC1 equivalent concentration) are between these two types of deposits. This kind of difference is observed in epithermal deposits in other countries (Hedenquist and Henley, 1985).
1.4.3.2. Estimate of temperatures from the electrum-sphalerite-pyrite-argentite assemblage As noted already, the Ag content of electrum in equilibrium with argentite and FeS content of sphalerite in equilibrium with pyrite are expressed as a function of fs2 and temperature, we can estimate temperature from Ag content of electrum and FeS content of sphalerite.
Miocene-Pliocene Hydrothermal Ore Deposits
125
Filling Temperature ('C) Au-Ag Dep.
150 [
200 i
250 ]
I
•
Sanru Teine-Takinosawa Todoroki-Chuetsu Todoroki-Shuetsu Chitose-Daikoku Chitose-Daikoku No2 Chitose-Fukujin Yatani-Kanizawa Takatama Sado-Ohdachi Nebazawa Seigoshi Taio Fuke-Honpi Ohkuchi Hishikari Arakawa Kushikino
I
3OO ] ]
•
•
350
I I
I
•
I
•
I
•
¢
I
: •
I
Pb-Zn'Mn Dep. Ohe-Senzai Inakuraishi Toyoha-Tajima Toyoha-Hadma Toyoha-lzumo Yagumo-Ohgid
•
I •
I
6
Cu" Pb oZn Dep. Oppu Osarizawa AnMnari Hosokura-Shoko Hosokura-Ohtake Hosokura-Hakuho Ohizumi Yatani-Honpi Nanetsu Tochigi Ogoya Taishu-Okutomi Taishu-Shinotomi Taishu-Himi Taishu-Taisho Taishu-Misoge Taishu-Tsurue Taishu-Akushidani Taishu-Shirocake Taishu-Amanohara
•
I
¢
¢
•
I
I •
]
I [
•
]
I
Figure 1.87. Summary of filling temperatures of fluid inclusions from Neogene vein-type deposits in Japan. Solid circIe represents average filling temperatures of fluid inclusions for individual deposits (Shikazono, 1985b).
Chapter 1
126 350
300
_~ 250
g
200
J1 ~
'I l
150
100 150
200 250 300 350 Electrum-Sphalerite Temperature (°(3)
400
Figure 1.88. Electrum-sphalerite temperatures vs. homogenization temperatures of fluid inclusions from epithermal Au-Ag vein-type deposits in Japan. Electrum-sphalerite temperatures were calculated from the iron content of sphalerite and the silver content of electrum (Equation 6 in Shikazono (1985a)). The average and range of homogenization temperatures and electrum-sphalerite temperatures for a given ore deposit are represented by a solid circle and line, respectively. 1 = Ohmidani-Fusei, 2 = Todoroki-Chuetsu, 3 = TodorokiShuetsu, 4 = Taio no. 9, 5 = Toyoha-Tajima, 6 = Nawaji, 7 = Seigoshi no. 2, 8 = Toi, 9 = Yugashima, 10 = Yatani-Tengu, 1l = Yatani-Kanizawa, 12 = Oh~Senzai, 13 = Ohe-Senzai, 14 = Sado, 15 = NebazawaManzai no. 3, 16 = Kamioka, 17 = Chitose-Daikoku, 18 = Yatani Honpi, 19 = Yunoura (Shikazono, 1985a).
In order to obtain reliable formation temperatures based on electrum-sphalerite geothermometer, the following conditions have to be satisfied. (1) The coexisting sphalerite, pyrite, electrum, and argentite must have been at e q u i l i b r i u m at the time of their precipitation. Although it is difficult to evaluate this condition, it is c o m m o n l y observed that these minerals are in direct contact with each other without evidence o f mutual replacement texture. Therefore, it is likely that these minerals have been precipitated nearly contemporaneously. (2) The FeS content o f sphalerite and the Ag content of electrum have not c h a n g e d considerably during the post-depositional period. It is unlikely that the sphalerite c o m p o s i t i o n c h a n g e d d u r i n g the cooling stage, because this mineral is one of the most refractory sulfide minerals. In general, the Ag c o n t e n t of electrum with a large grain size increases from core to margin. Such regularity o f compositional z o n i n g observed in a large grain, wide c o m p o s i t i o n a l variation in a large grain, and the relationship between grain size and compositional range in a grain suggest that the electrum composition is retained d u r i n g the post-depositional period. The A u - A g interdiffusion coefficient for electrum has b e e n experimentally studied ( C z a m a n s k e et al., 1973). The values of the A u - A g interdiffusion coefficients, the compositional z o n i n g pattern in electrum, and the temperature range of c o n c e r n (ca. 2 0 0 - 3 0 0 ° C ) suggest that the post-depositional change
M i o c e n e - P l i o c e n e Hydrothermal Ore Deposits
127
-5
-10
Y~
t~
0 _..1
-15
i
v
I
I
t
i
I
200
!
250
i
i
i
Temp.(°C)
I
i
300
Figure 1.89. Activity of S2(as~)-temperature diagram showing possible as2 and temperature ranges for epithermal Au disseminated-type (hot spring type), epithermaI Au-Ag vein-type and epithermaI base metal vein-type deposits in Japan (Shikazono 1986; Shikazono and Shimizu, 1988b). TABLE 1.13 Filling temperature and NaCI eq. concentration of fluid inclusions from epithermaI gold-silver and base-metal vein-type deposits (Shikazono and Shimizu, 1992) Deposit
Filling temprerature (°C)
NaC1 eq. concentration (wt%)
Gold-silver type
Sado Seigoshi Yatani Ohguchi Todoroki Koryu Chitose Kushikino
305-190 243-178 273-209 265-184 24(~ 122 300-140 300-220 250-210
2.5-1.0 2.8~0.0 1.5~).5 1.64).0 1.7-0.4 1.4-0.0 2.0-1.0 1.1~0.6
Base-metal type
Toyoha Oppu Osarizawa Hosokura Nanetsu Taisyu Asahi Ani Oe Jokoku
300-150 330-170 268-156 231 - 130 306-250 376-150 275 250 266-207 310-145 250-125
4.24).2 i8.3-1.7 7.5~0.0 9.1 ~).0 7.9-4.4 33.5~.0 10.3-8.8 12.0 2.0-0.0 6.0-3.0
128
Chapter 1
in electrum composition may be negligible. One of the most likely explanations for the compositional zoning in a grain is that it reflects the change in chemical parameters of ore fluids during the precipitation of electrum. (3) The value of the activity coefficients of FeS in sphalerite determined for temperatures above 300°C can be extrapolated to lower temperatures. As stated by Barton and Toulmin (1966), ?/FeS does not depend on temperature above about 270°C. However, the activity coefficient below 270°C has not been studied. Scott and Kissin (1973) have stated that activity coefficients for FeS in sphalerite at low temperatures may be substantially different from those at higher temperatures. (4) The effects of impurities such as Mn and Cd in sphalerite on the equations are negligible. Generally, the concentrations of minor elements in sphalerite from the epithermal A u - A g vein-type deposits in Japan, except for iron, are very small (Cd less than 1 wt%, Mn less than n 1 x 10 -1 w t % , Cu less than n x 10 -1 wt%, the concentrations of other elements are also less than n x 10 .2 wt%; e.g., Shikazono, 1978b). These low concentrations affect the fs2-temperature relations in the F e - Z n - S system (e.g., Barton and Toulmin, 1966). Impurities in electrum such as copper and antimony are also very small, generally less than 1 wt%. Therefore, it is likely that these elements do not affect the thermochemical properties of electrum. Argentite from epithermal Au-Ag vein-type deposits in Japan contains sometimes up to 10 wt% selenium (Shikazono, 1978b). However, if the selenium contents are small, it is likely that the activity coefficient for Ag2Se in argentite does not deviate significantly from unity, because complete solid solution between Ag2Se and Ag2S exists above ca. 180°C (Sugaki et al., 1982); However, the absolute value of the activity coefficient for Ag2S in argentite cannot be determined. Impurities in pyrite such as nickel and cobalt are also very low (less than n x 10 -1 wt%), and thus, an activity coefficient of FeS2 in pyrite equal to one can be safely assumed. (5) The effect of pressure is negligible. These epithermal A u - A g vein-type deposits have formed in a shallow and low-pressure environment and pressure correction, such that any correction to the homogenization temperatures will be small (probably less than 20°C). The temperatures estimated from the electrum-sphalerite-pyrite-argentite assemblage are plotted versus homogenization temperatures of fluid inclusions as shown in Fig. 1.88. Although the electrum-sphalerite temperatures and homogenization temperatures from given veins for each deposit have some variation, the averages of the electrum-sphalerite temperatures show a good correlation with the average homogenization temperatures. Most electrum-sphalerite average temperatures correlate to within 30°C of the respective average homogenization temperatures. This good correlation can be seen in Fig. 1.88. The good correlation between homogenization temperatures and electrumsphalerite temperatures suggests several points: (1) the uncertainties of electrumsphalerite temperatures are less than 20 ° to 30°C, even at temperatures from ca. 180 ° to 300°C, (2) the electrum-sphalerite-pyrite-argentite assemblage was formed close to equilibrium in Japanese epithermal Au-Ag vein-type deposits, and (3) the pressure corrections to homogenization temperatures for Japanese epithermal Au-Ag vein-type deposits is small, less than 20°C to 30°C. t All n in this text book as natural number 1 through 9.
Miocene-Pliocene Hydrothermal Ore Deposits
129
1.4.3.3. Gas fugacities
Sulflirfugacity (fs2)
As will be mentioned in section 2.4.3, i s 2 c a n be estimated based on the Ag content of electrum coexisting with argentite (or acanthite), the FeS content of sphalerite coexisting with pyrite and temperature estimated from homogenization temperatures of fluid inclusions. Figures 1.68 and 1.69 show the FeS content of sphalerite and the Ag content of electrum from epithermal Au-Ag vein-type, epithermal base-metal vein-type and Kuroko deposits, indicating different fs2-temperature regions for these types of deposition. Figure 1.89 shows typical range of fs2 and temperature for epithermal basemetal vein-type and Au-Ag vein-type deposits. It is noteworthy that the ranges of fs2 for epithermal Au-Ag, epithermal Au-bearing base-metal, and epithermal Au-free base-metal vein-type deposits are different, while temperatures are not different. As mentioned already, small amounts of electrum occur in epithermal base-metal vein-type deposits. Electrum is not observed in the epithermal base-metal vein-type deposits in which pyrrhotite occurs (e.g., Toyoha-Soya, Oizumi, and Hosokukura Pb-Zn deposits). However, electrum is found in epithermal base-metal vein-type deposits in which hematite is commonly observed (e.g., Osarizawa and Ani C u - P b - Z n deposits). This indicates that electrum precipitates in relatively high fs2 and f02 condition.
Oxygenfugacity (fo2). The fo2-pH diagrams (Figs. 1.90 and 1.91) were constructed at 200°C and 250°C based on the homogenization temperatures and electrum-sphalerite temperatures (Shikazono, 1985d). The ionic strength and activity coefficients of aqueous species are estimated from the freezing temperature of fluid inclusions. The ionic strength is assumed to be 1. Estimates of the total dissolved sulfur concentrations of ore fluids (ZS) responsible for several veintype deposits in Japan are of the order of 10-2-10 -3 mol/kg H20 (Shikazono, 1974b; Hattori, 1975). Chemical analyses of the hot springs accompanied by epithermal basemetal depositions (White, 1967; Browne and Ellis, 1970; Mayhon and Finlayson, 1972; Weissberg et al., 1979) give the values of 10 2-10-3 mol/kg H20. Therefore, the total dissolved sulfur concentration (ZS) was assumed to be 10-2-10 -3 mol/kg H20 for the construction of the diagrams. Gangue minerals and salinity give constraints on the pH range. The thermochemical stability field of adularia, sericite and kaolinite depends on temperature, ionic strength, pH and potassium ion concentration of the aqueous phase. The potassium ion concentration is estimated from the empirical relation of Na+/K + obtained from analyses of geothermal waters (White, 1965; Ellis, 1969; Fournier and Truesdell, 1973), experimental data on rock-water interactions (e.g., Mottl and Holland, 1978) and assuming that salinity of inclusion fluids is equal to mNa+ + mK+ in which m is molal concentration. From these data potassium ion concentration was assumed to be 0.1 and 0.2 mol/kg H20 for 200°C and 250°C. By giving the values of temperature, ZS, ionic strength, FeS content of sphalerite and Ag content of electrum, we can place a limit on fo2. However, we cannot know whether fo2 lies in the predominance field of reduced sulfur species or that of oxidized sulfur species from the constraints mentioned above.
Chapter 1
130
log fo 2
-35
i I
"".,,..
!
Ka ~, Se IAd -~,~,~_ i i Hm ,.~-r ---4 Mt I
B n,, Py I
~"~:L~."~-,~~ /
..A ~g - .
)
.
,
PY . . . . .-.'.- - - F
:%sit
Po
,
,
l i
I
I
-o"~. ~"
o,
, l
I
|
I
I
I
!
/
I
I
3
4
5
6
?
8
9
10
pH
Figure 1.90. fo2-pH diagram constructed for temperature = 200°C, ionic strength = 1, and ]ES = 0.01 moI/kg H20. Solid lines (1), (2), (3), and (4) are, respectively, (1) 10 tool% FeS in sphalerite; (2) 70 tool% Ag in electrum, (3) 60 mo1% Ag in electrum, and (4) 1 tool% FeS of sphalerite; 0.1 tool% FeS sphalerite is nearly coincident with the Bn + Py/Cp boundary. The dotted and shaded areas A and B represent the possible foz-pH regions for Au-Ag vein-type (Yatani, Seigoshi, Omidani, Asahi) and Pb Zn vein-type (Toyoha, Yatani, Ikuno) deposition, respectively. Ka: kaolinite, Se: sericite, Ad: adularia, Hm: hematite, Mt: magnetite, Bn: bornite, Py: pyrite, Cp: chalcopyrite, Arg: argentite, Nsil: native silver, Po: pyrrhotite (Shikazono, 1978b). The fo2 o f ore fluids responsible for the epithermal base-metal veins might have been in the predominance field o f reduced sulfur species because ( l ) pyrrhotite is occasionally found in these deposits, (2) selenium content of argentite is very low and (3) H2S is dominant in the present-day epithermal base-metal fluids. Implication of selenium content of sulfides will be considered later. Barite is sometimes found in the late-stage of mineralization. Thus, it is likely that fo2 of barite stage lies in the predominance field of oxidized sulfur species. As already discussed, fo2 o f Kuroko ore fluids is considered to lie in the predominance field of reduced sulfur species from the following two reasons; (1) Selenium content of sulfides is very low (Yamamoto, 1974) and (2) H2S is dominant in hydrothermal solution venting from back-arc basins (section 2.3) from which hydrothermal ore deposits being similar to Kuroko deposits form. On the other hand, the ore fluids responsible for epithermal A u - A g vein-type deposits contain appreciable amounts of oxidized sulfur species, together with reduced sulfur species. Oxidized sulfur species/reduced sulfur species ratio is considered to be greater than 1, that is, f o 2 lies in the predominance field of oxidized sulfur species. The reasons for this estimation are: (1) hematite is c o m m o n in the deposits, (2) barite is found
Miocene-Pliocene Hydrothermal Ore Deposits
131
log Io 2 -30
Ka] Se i Ad
~
|
~"'-
'
Hm Mt
~ + ~ - ? - 4 5 - , , , B n + P y "~,
.-'-(3)_ -(2)
CpN ~
--(I) - 4 0 ;__Arg____~' .
.
I PY
.
.
Nsi
.
,
=
\
Po
I
, I I
I
3
I I 11
4
l
\
I
I
I
I
I
f
5
6
7
8
9
10
pH
Figure 1.91. fo2-pH diagram constructed for temperature = 250°C, ionic strength = 1, and lgS = 0.01 mol/kg H20. Is~FeS mole percent lines for sphalerite and the stability relations of some hydrothermaI minerals are also given. The solid lines (1), (2), and (3) are respectively 10, 1, and 0.! mol% FeS in sphalerite; 70 and 60 mol% lines of Ag in electrum are nearly identical to lines (i) and (2). The dotted and shaded areas of A and B represent the possible fo2-pH regions for Au-Ag vein4ype (Yatani, Seigoshi, Omidani, Asahi) and Pb-Zn vein-type (Toyoha, Yatani, Ikuno) deposition, respectively. Bn: bornite, Py: pyrite, Cp: chalcopyrite, Arg: argentite, Nsi: native silver, Hm: hematite, Mr: magnetite, Po: pyrrhotite, Ka: kaolinite, Se: sericite, Ad: adularia (Shikazono, 1978b). in Te-type deposits, although this mineral formed at late stage and not found in Se-type deposits, (3) selenium is usually contained in the deposits, (4) tellurium is concentrated in Te-type deposits, and (5) sulfate ion is abundant, together with H2S in geothermal waters, associated with epithermal A u - A g depositions (section 2.1). Selenium and tellurium contents of sulfides are useful indicators for estimating f Q (Shikazono, 1974a, 1978b). The selenium content o f sulfide is governed by the reaction, (MS) + Se 2 - = (MSe) + S 2 -
(1-37)
where (MS) and (MSe) are MS and MSe components in the sulfide-selenide solid solution M ( S I - x , Sex), respectively. Equilibrium constant o f this reaction (K1-37) is, K1-37 = ( a M S e a s 2 ) / ( a M s a s e 2
).
(1-38)
where a is activity. Therefore, if the activity coefficient ratio, gMSe/gMS, does not deviate from unity and the temperature is constant, the ratio, mMse/mMS, correlates to ase2-~as2- which is represented as functions of f02, pH, temperature, and l ~ S e / 2 S (Shikazono, 1978b).
132
Chapter 1
Log (asga~-)
-1 -2 -3
1
-4
!
-5 -6
I
,
A
I
-7
J
-8 -9
I
I
i
(~)
!(2)
(3)
(4)
I(5) I
I
! i
i
i
11
i
7 2345678 Figure 1.92. Dependence of the ratio aSe2 ~as2
i
|
i
,
,i
i
i
~
,
91011121314
,
pH
pH under the conditions: temperature = 150°C, ionic strength = 1, li~S = 0.01 mol/kg H20, ~2Se = 10-7 mol/kg H20, and ZS = ]~rnr (total reduced sulfur content). (1) H2S-H2Se region, (2) H2S-HSe region, (3) HS HSe region, (4) S2 -HSe- region and (5) S2--Se 2 region (Shikazono, 1978b). on
Figure 1.92 represents the pH dependence of ase2 /as2- in the predominance region of reduced sulfur species. Considering the equation mentioned above, it is thought that this activity ratio correlates well with the selenium content of sulfide, if the effects of the activity of the MSe and MS components in sulfide (M: metal element) are neglected. The activity ratio, ase2 ~as2 , is expected to be relatively high in the sericite region and low in the adularia region. However, the selenium content of acanthite from epithermal A u - A g vein-type deposits in which adularia is a common constituent is higher than that from the epithermal Pb-Zn vein-type deposits in which sericite is the common wall-rock alteration product. Therefore, the difference in the selenium content of acanthite from these deposits cannot be explained by the difference in pH. As the values of the activity ratio, ase2 ~as2 , on the f o 2 - p H diagrams at 150°C (Fig. 1.93) and 300°C (Fig. 1.94) are nearly identical, changes in temperature alone cannot be responsible for the variation in the selenium content, if the temperature dependence of the equilibrium constant, KI-37, is not large. In contrast to the effects of pH and temperature, fo2 has a great effect on the selenium content of sulfides, especially in the predominance region of oxidized sulfur species (Figs. 1.93 and 1.94), where the activity ratio is relatively high and increases rapidly with increasing fo2 at a fixed pH. In the predominance region of reduced sulfur species, the ratio is constant at a fixed pH and has a relatively low value. The wide variation in the selenium content of acanthite from epithermal A u Ag vein-type deposits can be explained by assuming that the acanthite formed within the predominance region of oxidized sulfur species at a constant I2Se/~2S ratio, ionic strength, temperature and pH (Figs. 1.93 and 1.94). The very low selenium content of acanthite from epithermal Pb-Zn vein-type deposits suggests the predominance region of reduced sulfur species (Figs. 1.93 and 1.94).
Miocene-Pliocene Hydrothermal Ore Deposits
133
Io Ifo2 -30
-35 HSO~ (Na, K)SO4
-40 _ ~ -4s
__
~50
H2S
"-'
',HS~ S~-
. . . . . . . . . .
2-
H2SIHS i
i
[
HS S; ;I
i
1 2 3 4 5 6 7 8 9 1011 121314
pH
Figure 1.93. foz-pH diagram with the stability fields of aqueous species in Na-K-H-S Se-O system for the conditions: ]ES = 10 -2 moI/kg H20, ~2Se = 10 7 mol/kg H20, ionic strength = 1, and temperature = 150°C. Dashed lines are the ratio iso-ase2 /as2 in logarithmic units. Stability fields for native sulfur and native selenium and the boundaries between predominance regions of oxidized and reduced selenium species are omitted for clarity (Shikazono, 1978b).
Iogfo2
-20 -25
HSO~ (Na, K)SO~
-30 -35 -40
[
/-ii I--
-45
.
.
_0o
- -,
HS:; f
r
=
i
1 2 3 4 5 6 7 8 9 1011 121314
pH
Figure 1.94. fo2-pH diagram with the stability fields of aqueous species in the N a - K - H - S - S e - O system for the conditions: NS = 10 2 mol/kg H20, ]ESe = 10 7 mol/kg H20, ionic strength = 1, and temperature = 300°C. Dashed lines indicate iso-ase2 ~as2 contours in logarithmic units (Shikazono, 1978b).
Chapter 1
134
Log
a Carb
ZnCO3
1 i
O. -2' -4" -6' -8" -10,
U (2) -5o
-~
(3) t 6g fo 2
Figure 1.95. Activity of component ZnCO3 versus fo2- Carbonate containing ZnCO3 is in equilibrium with sphalerite. Thermochemical calculation was made under the following conditions: temperature = 200°C, ionic strength = 1, E;S = 10 2 m, and pH = 5. (1) CH4 and H2S region. (2) H2CO3 and H2S region. (3) H2CO3 and (Na, K) SO]- region (Shikazono, 1977b). In contrast to the estimation of total dissolved sulfur concentration (ES), the total dissolved selenium concentration (ESe) has not yet been estimated. The Beppu hot springs which are accompanied by A u - A g siliceous sinter, contain about 10 -7 tool/1 Se (Uzumasa, 1965). For the construction of Figs. 1.93 and 1.94, the Se/S ratio of the ore fluids is assumed to be 10 -5 . The difference in selenium content of acanthite from the different types of ore deposits can, of course, also be explained by the difference in Z S e / E S in the ore fluids; i.e., the ore fluids responsible for the formation o f epithermal A u - A g vein-type deposits may have had a higher E S e / E S ratio than that for epithermal P b - Z n vein-type deposits. It is likely that considerable amounts of sulfur were derived from marine rocks (Green tuff) and were incorporated into ore fluids for base-metal veins. The zinc content of siderite was applied to estimate fo2 of ore fluid by Shikazono (1977b). The analytical data on the late-stage siderites coexisting with barite, pyrite and hematite from the Ohmori epithermal A u - A g vein-type deposits that zinc content is high in the range of 0.8-5.8 wt% as ZnO, but the siderite of the early stage of mineralization from the Ohmori and from the Toyoha epithermal P b - Z n - A g vein-type deposits does not contain such a large amount of zinc. The relation between zinc content in carbonate and many physicochemical variables such as f % , temperature, pH, and so on was derived on the basis of the equilibrium between coexisting carbonate and sphalerite (Figs. 1.95 and 1.96). It is predicted theoretically that zinc content increases with increasing fo2 in the oxidized sulfur species and oxidized carbon species region (Fig. 1.95). This estimate is consistent with mineral assemblage containing the siderite. The iron content of sphalerites coexisting with siderite and pyrite from the Toyoha deposits is high (3-12 wt%). On the other hand, sphalerite of later stage Ohmori
Miocene-Pliocene Hydrothermal Ore Deposits
135
Log fo 2 -35 HSO41(Na,K)SO 4
HCO~ CO~-
H2C03 HC03
I
+F;; • p ;..=,~~ . ,.,-:..:. . . . . . . . . . . . ... .t. . . . . . .
;
-40
. . . ~ . : ~ - - -- 2(2)
/
Bn+P, ~ "L ""~; ' ~ ' ' " (3) | ..................... ,,. . . . ~ '.':..( Na, K)S02 Cp H~CO~ ! i I I "~ ~ Z-45 D,,
"I
--
: : I, ' .....
1
"°
-5o
i
i
Ka Se Se:A i
4
~.-.aE~!Bn 2 ',~.-... CO.C ',,1.'>E-..J ~I ;! ~"C HL4 T M Po',M t
H2s HS- Hg S2-
t 2
} ] CH4 : I -"~. :j I i, J._,. -~ ~" '
5
6
7
\
\
! g
lb
pH
Figure 1.96. Log foz~H diagram constructed for temperature = 200°C, ionic strength = 1, }IS 10 -2 m, and P,C = 10 tm. Solid line represents aqueous sulfur and carbon species boundaries which are loci of equal molalities. Dashed lines represent the stability boundaries for some minerals. Ad: adularia, Bn: bornite, Cp: chalcopyrite, Ht: hematite, Ka: kaolinite, Mt: magnetite, Po: pyrrhotite, Py: pyrite, Se: sericite. Heavy dashed lines (1), (2), and (3) are iso-activity lines for ZnCO3 component in carbonate in equilibrium with sphalerite: (1 "~ a carb = 0 . 1 . (2) carb carb -(Shikazono,1977b). J ZnCO3 azncO3= 0 . 0 1 . (3) azncO 3-0.00I =
deposits contains small amounts of iron (0.1-0.4 wt%), and the content is narrow in range (although the iron content in early-stage is higher; 0.4-6.2 wt%). The decrease o f iron content from the early to the later stage indicates that oxidation occurred a n d / o r temperature decreased during the formation of the Ohmori deposits. The difference in iron content for these deposits means that the Toyoha deposits formed under relatively lower fo2 conditions than the early stage Ohmori deposits. These conclusions derived from the analytical data on coexisting siderite and sphalerite are consistent with the studies on the chemical environment of epithermal vein-type deposits in Japan (Shikazono, 1973, 1974a; Hattori, 1975).
(fc02), The fCO2 values can be estimated from (1) gangue mineral assemblages including carbonates and (2) fluid inclusion analyses. Shikazono (1985b) summarized the assemblage and mode of occurrence o f common gangue minerals from more than 70 Neogene epithermal vein-type deposits in Japan. Carbon dioxidefugacity
Chapter 1
136
3
J
2 1
~ 0 o M , , ~ k , , ~_
/i
.,d
-2 -3 -4
150
200
250
300 Temperature (°C)
350
Figure 1.97. log fco2-temperature diagram showing the univariant equilibrium curves for some gangue minerals. A. 2Ca2AI3Si3012(OH) (clinozoisite) + 3 SiO2 (quartz) + 2CACO3 (calcite) + 2 H20 = 3 Ca2Al2 Si3OI0(OH)2 (prehnite) + 2CO2 (Xpis = 0.30, Xpis: mole fraction of pistacite component in epidote). B. Ca6Si60]7(OH)2 (xonotlite) + 6CO2 = 6CACO3 (calcite) + 6SIO2 (quartz) + H20. C. CaCO3 (calcite) q- TiO2 (rutile) + SiO2 (quartz) = CaTiSiO5 (sphene) + CO2. D. MnSiO3 (rhodonite) + CO2 = MnCO3 (rhodochrosite) + SiO2 (quartz). E. 3KA13Si3OIo(OH)2 (K-mica) + 4CACO3 (calcite) + SiO2 (qua~z) = 2Ca2A13Si3OI2(OH) (clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 4- 2H20 (Xpis = 0.3). P. 3KAI3Si3Om(OH)2 (K-mica) + 4CACO3 (calcite) + SiO2 (quartz) = 2Ca2AI3Si3OI2(OH) (clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 Jr- 2H20 (Xpis = 0.25). G. 3FeCO3 (siderite) + (1/2) 02 = Pe304 (magnetite) + 3CO2, C (graphite) + 02 = CO2. H. 3CaMg(CO3)2 (dolomite) + KAISi308 (K-feldspar) 4- H20 = 3CACO3 (calcite) + 3CO2 4- KMg3(AISi3Om)(OH)2 (phlogopite). I. C (graphite) + 02 = CO2, FeS (pyrrbotite) + 1/2S2 = FeS2 (pyrite), 2H2S(aq) + 02 = $2 4- 2H20(I). J. FeCO3 (siderite) + Fe203 (hematite) = Fe304 (magnetite) + CO2. K. CaA12Si4OI2-2H20 (wairakite) + KAISi308 (K-feldspar) + CO2 = CaCO3 (caIcite) + KAI3Si3Oto(OH)2 (K-mica) + 4SIO2 (quartz). L. CaAI2Si4OI22H20 (wairakite) + CO2 = CaCO3 (calcite) + AI2Si2Os(OH)4 (kaolinite) + 2SIO2 (quartz). M. CaAI2Si4OI2-4H20 (laumontite) + CO2 = CaCO3 (calcite) + AI2Si205(OH)4 (kaolinite) + 2 SiO2 (quartz) + 2 H20 (Shikazono, 1985b).
As already mentioned the major gangue minerals vary with different deposit types; quartz, chalcedonic quartz, adularia, calcite, smectite, interstratified mica/smectite, interstratified chlorite/smectite, sericite, zeolites, and kaolinite in Au-Ag deposits, chlorite, quartz, sericite, calcite, rhodochrosite, siderite and (magnetite) 2 in Pb-Zn-rich deposits, chlorite, sericite, siderite, hematite, magnetite and (epidote) in Cu-rich deposits. Based on the gangue mineral assemblage (Fig. 1.97), homogenization temperatures of fluid inclusions, thermochemical calculations (Fig. 1.98), and analytical data on fluid inclusions (Fig. 1.99), typical ranges of fco2 for the Au-Ag, Pb-Zn-Mn and Cu-Pb-Zn vein-type deposits were determined to be 10 .3 to 1 atm (190°C-250°C) l0 -I to 10 atm (200-250°C) and 10- I to 103 atm (200°C-350°C), respectively (Fig. 1.100). The fco2 of Kuroko ore fluids is close to that of Cu-Pb-Zn vein-type deposits. The fco2 values could be estimated based on FeCO3 content of carbonates coexisting with iron minerals (pyrite, hematite, magnetite, pyrrhotite) and minerals containing 2 Mineral in parentheses occurs in small amounts in each deposit type.
Miocene-Pliocene Hydrothermal Ore Deposits
137
3 2
1
# o
0
~-1
._1
-2 -3 -4
150
200
250
300 Temperature (°C)
350
Figure 1.98. Summary of fcQ-temperature ranges for Au-Ag-rich, Pb-Zn Mn-rich, and Cu-Pb-Zn-rich vein-type deposits based on gangue mineral assemblages and fluid inclusion data. Line a-b: Equilibrium among graphite, pyrite, and pyrrhorite. Line c~l: fc02 vs. temperature curve for fluid containing i tool% CO2 with temperature of first boiling of 300°C. Line e-f: "Plagioclase" + CO2 = calcite + "kaolinite". Line g ~ : fco2 temperature relation for geothermal reservoir waters obtained by Arn6rsson (1984). Line i-j: f c Q vs. temperature curve for fluid containing 1 tool% CO2 with a temperature of first boiling of 350°C. Line ~1: 2Ca2A13Si3012(OH) + 3KAISi308 + 4CO2 + 2H20 = 4CACO3 + 3KAI3Si3OI0(OH)2 + 6SIO2; mole fraction of KAI3Si3010(OH)2 in mica = 0.6, mole fraction of 2Ca2Fe3Si30[2(OH) in epidote = 0.2 (Shikazono, 1985b). 31
yA
2! 1 S
04
O o
0
C
P
o
_.1 -2 -3 -4
1;0
2;o
2;0
330 T e m p e r a t u r e (°C)
Figure 1.99. Estimated fco2-temperature ranges from anaytical data on fluid inclusions and homogenization temperatures (Shikazono, 1986). T: Taishu (Pb, Zn), O: Ohizumi (Cu, Pb, Zn), Y: Yatani (Pb, Zn), Os: Osarizawa (Cu, Pb, Zn), H: Hosokura (Pb, Zn), C: Chitose (Au, Ag), S: Seigoshi (Au, Ag).
iron as solid s o l u t i o n ( i r o n c o n t e n t in s p h a l e r i t e ) . S h i k a z o n o ( 1 9 7 4 a ) t h e o r e t i c a l l y d e r i v e d the r e l a t i o n s h i p b e t w e e n iron c o n t e n t o f c a r b o n a t e s c o e x i s t i n g w i t h i r o n m i n e r a l s as f u n c t i o n s o f f c Q , f Q , a n d others. W e c o u l d e s t i m a t e f c Q u s i n g his m e t h o d . B u t n o q u a n t i t a t i v e a p p l i c a t i o n h a s b e e n c a r r i e d out.
Chapter i
138
3 2
1
020 o,-1 -2 -3 -4
150
200
250
300 350 Temperature (°C)
Figure 1.100. Typical fco2-temperature ranges for Au-Ag-rich, Pb-Zn-Mn-rich, and Cu-Pb-Zn-rich veintype deposits estimated from gangue mineral assemblages, homogenization temperatures of fluid inclusions, and thermochemical calculations (Shikazono, 1985b).
Seleniumfugacity (fse2).
If coexisting electrum, sphalerite, pyrite, argentite and galena are in equilibrium, the relationship between the Ag content of electrum, selenium contents of argentite and galena, iron content of sphalerite, temperature, fsz and fse2 can be derived from the equilibrium relations for the chemical reactions given in Table 1.14. Scott and Barnes (1971) have obtained an equation representing FeS content of sphalerite in equilibrium with pyrite as functions of fsz and temperature. By combining equations (1)-(5) in Table 1.14 and the relation between activity coefficient of Ag in electrum, Ag
TABLE 1.14 Chemical reactions and equations representing the equilibrium relations used for drawing Fig. 1.101 Chemical reactions
Equilibrium relations
Temperature range (°C)
4 Ag + 82 (g) = 2 Ag2S (acanthite)
(I)
log f s 2 = ( - 9 7 9 0 . 2 I / T ) + 4.83
41ogaAg
(1)
25-176
4 Ag + $2 (g) = 2 Ag2S (argentite)
(2)
log fs2 = (--9173.95/T) + 3.61 + 2 log aAg2s - 4 log aAg
(2)
I76-804
4 A g + Se2 (g) = 2Ag2Se (naumannite) (3)
Iog fSe2 = (-- 10644.67/T) + 3.12 + 21ogaAg2sc 4IogaAg
(3)
133-727
+ 21ogaAg2s
4PbS + Se2 (g) = 2Pb2Se
(4)
log iS2 = log fSe2 q- 755.68 -- 0.24
FeS + 1/2S2 (g) = FeS2 (pyrite)
(5)
log fs2 = -- 15460/T + i4.32 - 2 log XFcS (5)
- 2 Iog(apbse/apbs)
(4)
25-327 ca. 400-700
T: absolute temperature; fs~: sulfur fugacity; fSe2: selenium fugacity; aAg2S: activity of Ag2S in argentite; aAg2Se: activity of Ag2Se in argentite; aAg: activity of Ag in electrum; apbse: activity of PbSe in galena; apbs: activity of PbS in galena; XFeS: FeS mole fraction of FeS in sphalerite (Shikazono and Takeuchi, 1984).
Miocene-Pliocene Hydrothermal Ore Deposits -10
139
-10
o=
8~
-15
~ 2 1
~:~ -15
o' o
~-"'~'-
~,--- 22
-20 H.T
-25 . . . . . . . 150
!. . . 200 . . . . . . . . . . 250 300 Temperature (°C)
-25
. . . . . . . 150 . . . . . . . . . . 200 ....
250
300
Temperature (°C)
Figure 1.101. Selenium fugacity-temperature diagram, i3: Argentite (or acanthite)-electmm-galena-Se2(g) equilibrium curve for XAg2S (X: mole fraction) = 0.8, Xpbse = 0.05 and XAg = 0.4. 14: Argentite (or acanthite)-electrum-S2(g) equilibrium curve for XAg2Se = 0.2 and XAg = 0.4. 15: Argentite (or acanthite)--electrum-galena-Se2(g) equilibrium curve for XAg2S= 0.8, XpbSe= 0.05, and XAg = 0.8. 16: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg2Se = 0.2 and XAg = 0.6. 17: Sphalerite-pyrite-galena-Se2(g) equilibrium curve for Xr,bSe= 0.05 and X F e S = 0.0l. 18: Sphalerite~yrite-galena-Se2(g) equilibriumcurve for Xpbse = 0.05 and XFeS= 0.02. 19:Sphalerite-pyrite-galena-Se2(g) equiiibriumcurve for Xpbse= 0.02 and XFeS= 0.004. 20:Sphalerite-pyrite-galena-Se2(g) equilibriumcurve for Xpbse = 0.02 and XFeS= 0.007. 21: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg = 0.5 and XAg2Se = 0.5. 22: Argentite (or acanthite)-galena-electrum-Se2(g) equiIibrium curve for Xpbse = 0.02, XAg2S~---0.5 and Xag = 0.5. 23: Argentite (or acanthite)-electrum-Se2(g) equilibrium curve for XAg2Se = 0.5 and XAg = 0.7. 24: Argentite (or acanthite)-galena--electrumSe2 (g) equilibrium curve for XAg2S= 0.5, XpbSe= 0.02 and XAg = 0.7. Estimated ranges of temperature and fSe2 for Kushikino and Takatama are shown in this figure as shaded areas. H.T.: homogenizationtemperature of fluid inclusions.
content of electrum and temperature obtained by White et al. (1957), the relationships between temperature, fs2, fse2, Ag content of electrum, FeS content of sphaterite, selenium contents of argentite (or acanthite), and of galena can be derived (Fig. 1.101). From these relations and analytical data on coexisting galena, sphalerite, argentite and electrum, the formation temperature, fs2, and fse2 can be estimated (Shikazono and Takeuchi, 1984) (Fig. 1.101). To derive the relations shown in Fig. 1.101, unity of activity coefficients of FeS2 in pyrite, Ag2S, and Ag2Se in argentite (or acanthite) was assumed. Activity coefficient of FeS2 in pyrite should be very close to unity, because the concentrations of minor elements (e.g., Ni, Co) in pyrite are very low, less than n x 10 . 2 wt%. Bethke and Barton (1971) have suggested in their experimental study on the distribution of selenium between coexisting galena and sphalerite that the P b S - P b S e system behaves as ideal solid solution at least above 600°C. Ag2S in argentite (or acanthite), FeS in sphalerite and Ag in electrum are 5 mol%, ca. 20 mol%, 1.2-2.4 mol%, and 4 6 - 6 0 atomic%, respectively, for the Kushikino deposits and 2 mol%, ca. 50 mol%, 0.4-0.7 mol%, and 5 0 - 7 0 atomic%, respectively for the Takadama deposits. Detailed descriptions on these deposits can be referred to in Kitami (1973), Sukeshita and Uemura (1976), Yamaoka and Nedachi (1978b), Izawa et al. (1981) and Takeuchi and Shikazono (1984). Based on these analytical data on the minerals mentioned above and thermochemical consideration formation temperature, fs2, and fse2 for these deposits are
140
Chapter 1
estimated. Temperature and fs2 estimated on the basis of fs2-temperature diagram is ca. 220-300°C and 10-9-10 -13 atm for the Kushikino and ca. 150-250°C and 10-11-10 -18 atm for the Takadama. Temperature and fse2 estimated on the basis of fSez-temperature diagram (Fig. 1.101) is ca. 200-300°C and 10-12-10 -18 atm for the Kushikino and ca. 150-250°C and 10-14-10 -23 atm for the Takadama. Homogenization temperatures of fluid inclusions in quartz for the electrumsphalerite-pyrite-argentite-galena stage of the Kushikino and Takadama deposits are in the range of ca. 180-250°C (Takeuchi, 1979; Izawa et al., 1981) and ca. 160-240°C (Yamaoka and Nedachi, 1978b; Watanabe, 1979), respectively. Wright et al. (1965) have found from hydrothermal experiments that continuous solid solution exists in the PbS-PbSe system at 300°C. Coleman (1959) has described specimens covering the entire range of solid solution between galena and clausthalite from vanadium-uranium deposits of the Colorado Plateau type. Thus, it seems likely that the departure from ideal solid solution in the PbS-PbSe system is not large in the temperature range considered here (ca. 180-300°C). However, unfortunately, absolute values of VPbS and YPbSe (Y: activity coefficient) in the temperature range considered here cannot be estimated. It is known that continuous Ag2S-Ag2Se solid solution exists in the temperature range concerned (ca. 180-300°C), although this solid solution is unquenchable. Deviation from ideality for AgzS-Ag2Se solid solution is also not studied. Analytical data on galena, argentite (or acanthite), electrum and sphalerite which all coexist with pyrite in the Se-rich Au-Ag vein-type deposits are available from Kushikino and Takadama Au-Ag vein-type deposits (Kawai, 1976; Takeuchi, 1979). PbSe in galena, Ag2Se in argentite (Takeuchi, 1979) and thermochemical calculations are used to estimate temperature and fs2 and Jse2- The temperature estimated from this mineral assemblage for the Kushikino seems to be slightly higher than the homogenization temperature of fluid inclusions. There are several possible reasons for this discrepancy. They are: (1) uncertainties of free energy changes for the reactions in Table 1.14, (2) changes of chemical compositions and phases during the post-depositions stage, (3) not simutataneous precipitations of electrum, sphalerite pyrite, galena, argentite (or acanthite) and quartz which are studied for the chemical composition and fluid inclusions, and (4) deviation from ideality of Ag2S-Ag2Se and PbS-PbSe solid solutions. It is difficult to determine which is the main cause for this discrepancy. In order to solve this, more detailed studies on this assemblage and thermochemical properties of PbS-PbSe and Ag2S-AgzSe solid solutions are required. Although such a discrepancy exists, it was for the Se-rich A u - A g vein-type deposits (Kushikino and Takadama) that (l) the mineral assemblage of sphaleriteelectrum-argentite-galena-pyrite assemblage is a useful indicator of environmental condition (fse~, fs2 and temperature), (2) precipitation temperature for this mineral assemblage is in the range of 150-300°C, and (3) fse2 is lower than fs2. Based on the .[Se2, fs2 and temperature estimated from this mineral assemblage, we can place a limit on the ranges of the other important chemical parameters such as total dissolved selenium and sulfur contents in ore forming solution responsible for the Se-rich A u - A g vein-type deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
141
1.4.3.4. Chemical composition of ore fluids The solubilities of ore metals depend on several variables such as fs2, fo2, pH, salinity, NS and temperature (Barnes and Czamanske, 1967; Shikazono, 1972b; Barnes, 1979). In the earlier sections, we estimated the ranges of these variables. Therefore, it is possible to calculate chemical compositions of ore constituent elements in ore fluids. For instance, Au concentration of ore fluids responsible for epithermal A u - A g deposits could be estimated from the following reaction. Au + HS + H2S + l / 4 02 ----Au(HS) 2 + 1/2 H20
(1-39)
Seward (1973) experimentally determined the solubility of Au due to this complex and equilibrium constant for the above reaction. Figure 1.102 shows the solubility of Au on l o g f o z - p H diagram calculated based on the thermochemical data by Seward (1973). The solubility of gold is high in neutral and near oxidized sulfur species/reduced sulfur species boundary (Fig. 1.102). It is noteworthy that this region corresponds to the f o z - p H region of epithermal A u - A g vein-type depositions (Figs. 1.90 and 1.91) (Shikazono, 1974a; Hattori, 1975). The solubility of pure gold is shown in Fig. 1.102. However, gold does not occur as pure gold in A u - A g deposits, but as electrum. Therefore, the effect of Ag content of electrum on gold solubility must be considered. Shikazono and Shimizu (1987) estimated NAu/NAg ratio in ore fluids responsible for Au-Ag veins based on the following reaction. (Au)e] + AgC12 + 2H2S = (Ag)el + Au(HS) 2 + 2C1- + 2H +
(1-40)
where (AU)el and (Ag)el are the Au and Ag components of electrum, respectively. Au
I HS021 S0z~
30
HEMATITE
250 °C m:~S = 3 x 10-3
I
40 PYRRHOTITE MAGNETITE
45
H2SI HS- \ 2
I
I
4
6 pH
i I
8
Figure 1.102. fo2-PH diagram at 250°C showing the stability fields of the principal sulfur species and solubiiity contours for gold in mg/kg as Au(HS)~-(HenIey,1984).
Chapter 1
142 From the equilibrium relation for equation (1-40), we obtain, aAg/aAu =
2 2 2 (au2smagcl~K1-40)/ (mau(us)~aci_aH+)
(1-41)
where K1-40 is equilibrium constant for the reaction (1-40), and ?/AgCl~/)/Au(rtS)~ is assumed to be unity. The equation is controlled by temperature, acl , aH2S, pH and mAgc12/mAu(HS)~ratio. We can calculate the E A u / E A g ratio in ore fluids responsible for Au-Ag veins, by giving temperature, salinity, pH and H2S concentration. The temperature is estimated based on fluid inclusion studies and electrum-sphalerite-argentite-pyrite assemblage. The NaC1 equivalent concentration of ore fluids is approximated from freezing temperature data on inclusion fluids, though the final melting temperature of fluid inclusion ice is also affected by CO2 concentration in epithermal ore fluids (Hedenquist and Henley, 1985). The pH values are estimated assuming the equilibrium among K-feldspar, K-mica, and quartz; this in turn allows a calculation of the potassium ion activity. The activity of HzS is estimated based on the equation showing the relation between the partial pressure of H2S gas and temperature of active geothermal waters (Giggenbach, 1980; Arn6rsson, 1985). Using the typical value of these variables and XAg (mole fraction of Ag in electrum) = 0.5, which is a typical value for electrum from epithermal Au-Ag vein-type deposits, E A u / E A g is calculated to be about 0.1, which is similar to that of Broadlands geothermal water (New Zealand) which is associated with gold precipitation. The above calculation is based on the assumption that AgC12 is the predominant Ag species in ore fluids. However, it is possible that silver bisulfide complex could contribute significantly to the transportation of Ag (Henley, 1985; Brown, 1986). If gold and silver bisulfide complexes are the dominant gold and silver aqueous species in Broadlands geothermal water and ore fluids responsible for Japanese epithermal Au-Ag vein-type deposits, the Ag/Au ratio of ore fluids responsible for Japanese Au-Ag epithermal vein-type deposits may be similar to that of Broadlands geothermal water which is about 0.1. tf mci- and pH are assumed to be 1-5 molal and lower than that for K-feldspar-Kmica-quartz equilibrium, respectively, E A u / E A g is estimated to be considerably lower than 0.1. Therefore, E A u / E A g of ore fluids for epithermal base-metal vein-type deposits is thought to be considerably lower than 0.1. The concentrations of base-metals (Cu, Fe, Pb, and Zn) in hydrothermal solution in equilibrium with sulfides (chalcopyrite, pyrite, galena and sphalerite) depend on several variables such as pH, mcl- concentration, temperature, mN2S, and fo2. The relation between the concentrations and these variables can be derived based on the chemical equilibrium for the following reactions. CuFeS2 + 2 C1- ÷ 2 H + + 1/2 02
~---FeS2 +
CuC12 ÷ H20
(1-42)
FeS2 + 2 C1- + 2 H + + H20 = FeCI2 + 2 H2S + 1/2 02
(1-43)
ZnS + 2C1- + 2H + = ZnCI2 + H2S
(1-44)
The ranges of these variables of epithermal ore fluids were estimated in section 1.4.3. Although the ranges of these variables are wide, we could estimate the concentrations of base metal elements if we took typical ranges of these variables.
Miocene-Pliocene Hydrothermal Ore Deposits
143
1.4.3.5. Stable isotopes 6D and 3180. 3D and 3180 of the ore fluids responsible for epithermal A u - A g and basemetal vein-type deposits in Japan have been estimated from analyses of fluid inclusions (Hattori and Sakai, 1979) and minerals (Watanabe et al., 1976). These data are shown in Fig. 1.103. 3D values of ore fluids for epithermal A u - A g vein-type deposits are similar to those of present-day meteoric water values. 3D values of epithermal ore fluids for base-metal vein-type deposits are slightly higher than those of epithermal A u - A g vein-type deposits. This may be due to the boiling of epithermal base-metal ore fluids and involvement of seawater. 3180 values of ore fluids are higher than those of meteoric water values. This is considered to be due to oxygen shift, which was caused by meteoric water-rock interaction. 3180 of ore fluids responsible for epithermal base-metal ore fluids is higher than that for epithermal A u - A g ore fluids. This is due to high extent of water-rock interaction and probably involvement of seawater and igneous water in the ore fluids. In individual deposit, 3180 of minerals varies widely. For example, 3180 of quartz and adularia increases with the stage of mineralization, and correlates to adularia/quartz ratio (A/Q), and Au and Ag grades of ore (Fig. 1.104). Such increase in 8180 is found also in carbonates in the Seigoshi A u - A g deposits and quartz and adularia in the Hishikari A u - A g deposits (Shikazono, 1988a; Shikazono and Nagayama, 1993). This increase is interpreted in terms of boiling of fluids at late-stage (Shikazono, 1989), decrease in water/rock ratio (Shikazono and Nagayama, 1993) and/or an influence of sedimentary rocks (Shikazono, 1999b).
~
.
-20-40t~
J
SG
S.W.
YUG
-60-80-
-100
I
-10
1
I
I
-8
-6
-4
I
I
-2
0
1
2
4
Figure 1.103.3D and 3180 of ore fluids responsibIe for epithermai Au-Ag vein-type deposits in Japan (Hattori and Sakai, I979: Imai et aI., 1998). S.W.: seawater, M.W. line: meteoric water line, KK: Kushikino, SG: Seigoshi, YUG: Yugashima, TK: Takatama, FUK: Fuke, YN: Yatani, KN: Kanisawa (Yatani), HK: Hishikari.
Chapter 1
144 ©
=
=
0
>
-1-
< 1" 0"
-2
-3. 0.5"
i "
B
-4
Pb
-5
L
O, -2.
-3.
• Quartz-Adularla ratio • 6180 Fluid
© ,.d
•
-4 3-1T I
1 3 6-IT
i 6 [
, 9 6-2T
i
/ 12 6-3T
~Au
I
i I5 64T
J I thickness 18 21 I 5-1T I 5-2T I (cm)
Figure 1.104. The relationship among 8180 (in permil) of fluids, minor element contents, and A/Q (adularia/quartz ratio) in the vein from wail rock side (3-IT) to central part (5-2T) (Ryosen No. 5 vein, Hishikari mine, 85 ml E50) (Shikazono and Nagayama, 1993).
8D and 8180 values for epithermal deposits from other countries are summarized in Fig. 1.105 (Field and Fifarek, 1985). The oxygen shift away from the meteoric water line is always observed, but 8D is similar to meteoric water value, suggesting meteoric water source of epithermal ore fluids. Magmatic contribution to ore fluids has not been found except in some ore fluids responsible for the deposits in the other countries; Tui
Miocene-Pliocene Hydrothermal Ore Deposits
145
8180 %0 -20 0
,
-10
0
i " /; Meteoric Water LinV
-60
Ocean Water iSMOW;
= /
//
&TF
Walrakel/Broadlandsl'.~ LarderelloQ~". . . . . Gov~,ar~/aAW C
Magmatic water 4 [ 1 &GS I•CP
Salton SeaZ.~- . . . . . . . .
-80 -100
+20
/
-20 -40
+10
I II
,/.~
~ao/%~
/
Mt. Lassen(~-----.-~ . . . . . . . II/~T .Uu ATE =11
_~/ oO~/e¢~
/,o ,G ,,~ . ~~/ : ,~q# =~_~--oSteamboat Springs /~-~ / / R&•A I " " CO& / / - " /
-120
6"g"
Geothermal H2' • Surface * Subsurface Hydrothermal H2' • Epithermal
//
-140 ;Y~owstone
/
/_® t____J ~,~ # /..~-
.-e
/ -
// /
CA~/'/
/
-160 I
i
I
f
I
I
I
I
I
Figure 1. i05. Distributions of 3D and 3]80 in various waters, minerals, and hydrothermal fluids of epitherma] deposits (FieId and Fifarek, 1985).
(New Zealand) (Robinson, 1974); Finlandia (Kamilli and Ohmoto, 1977), and possibly Comstock Lode (Taylor, 1973). ~13C and 3180 of carbonates. 313C and 3]80 of carbonates have been obtained by
Osaki (1973), Matsuhisa et al. (1985), Shikazono (1988a, 1989), and Morishita (1993) (Fig. 1.106). Roughly, the data lie between the igneous value (-7%0 to -8%o), and marine carbonate value (-1%o to +4%0). These are similar to the variations in carbonates from the Kuroko mine area. However, 313C and 3180 of carbonates from the epithermal vein-type deposits vary more widely than the carbonates from the Kuroko mine area. Some data plot in the region below the igneous carbonate-marine carbonate mixing line. This suggests that meteoric water was involved in the ore fluids. On the other hand, some data deviate from igneous carbonate-marine carbonate mixing line to higher 3180 region. This higher ~180 values are explained in terms of boiling of ore fluids. High 3180 values are obtained for carbonates of late-stage of mineralization (Shikazono, 1988a). Figure 1.107 shows the frequency of 313C of carbonates from epithermal A u Ag vein-type deposits and that from base-metal vein-type deposits. The carbonates are divided into two types: type A and type B. Type A is characterized by: (1) abundant occurrence in each deposit; (2) coexistence with sulfide minerals; and (3) large grain size. Main carbonate minerals are rhodochrosite and Mn calcite, whereas calcite is the main carbonate mineral for type B. Mn-carbonates of type A occur in P b - Z n - M n vein-type deposits. Type B is characterized by; (1) poor amounts in each deposit; (2) coexistence
146
Chapter 1
25 o o
o
2O
oe
og
o
oo
-to
o
o
o ee
o
o oo
•
o
.
o
• o
o g
g15
E O .IQ
e~
•
~dl0
o•
oS t ~ . ~ , , o ~ . " , % , " Oo°o o
-,..;:
o
~5
°o•
°o.**
Oo , *.~
t..~
o
o
o .
•
°o
o
o
ee
o
oe
| ~o
-5 $13C of C a r b o n a t e s ( ' / . , ) Figure 1.]06. 5180-513C of carbonates from Neogene vein-type deposits in Japan (open circle = calcite; solid circle = rhodochrosite and Mn-ca]cite; solid triangle = dolomite; cross = siderite) (Shikazono, ]989).
with late-stage quartz occurring in vugs; and (3) small grain size. Type B carbonate occurs in Au-Ag vein-type deposits and in Cu-Pb-Zn vein-type deposits but not in Pb-Zn-Mn vein-type deposits. The range of 313C of Japanese epithermal veins is similar to that of the other countries which is in a range of -10%o to 0%0 (Fig. 1.108 (Field and Fifarek, 1985). 313C and 3180 of carbonates from southern Kyushu (Hokusatsu gold district) have been studied in detail (Matsuhisa et al., 1985; Morishita, 1993). Morishita (1993) found that the 313C values of hydrothermal solution in the district during the mineralization stages were low (-11%o), compared with that of average crustal carbon (-7%o), suggesting that 313C of hydrothermal solution is controlled by organic carbon in widely distributed sedimentay rocks of the Cretaceous Shimanto Supergroup basement. 313C and 3]80 of carbonates in the Seigoshi mine district, Izu Peninsula, middle part of Honshu were determined by Shikazono (1988a). He showed that early-stage carbonates have 313C and 3180 values of -2.9%0 to +0.6%0 and +1.7%o to +10.2%o, respectively, suggesting a contribution of marine carbonate in Miocene marine sediments (Green tuff), but late-stage carbonates have high 3180 and low 313C, suggesting the effect
Miocene-Pliocene Hydrothermal Ore Deposits
147
25
20
z
5
1 -10
-8
-6
-4
-2
a'~c (%o)
(b)
25
> , 20 E
<
15
0
-Q 10
E z
5
8130
(%0)
Figure 1.107. Frequency histogram for ~ 13C of type-A(a) and type B (b) carbonates (Shikazono, 1989).
of boiling. Osaki (1973) suggested that the higher 3~3C (-2.5%0 to -7.5%0) and ~]80 values (+7.0%o to + 15.0%o) of rhodochrosites from the Jokoku Mn vein in Southern Hokkaido are due to contamination with limestone. In addition to boiling and origin of carbon, temperature and chemical states of dissolved carbon influence 313C of carbonates (Matsuhisa et al., 1985). These detailed studies on individual mine district suggest that carbon in carbonates was derived from the country rocks underlying the ore deposits and oxygen in ore fluids is controlled by origin of ore fluids (mostly meteoric water) and boiling of ore fluids.
~34S of sulfides. A large number of ~34Sdata on sulfides from epithermal base-metal and A u - A g vein-type deposits are available (e.g., Shikazono, 1987b). The 334S data are summarized in Fig. 1.109. The ~34S values of A u - A g deposits range from -7.5%0 to +5%o. However, majority of the ~34S values fall in a narrow range from -1%o to +3%e. ~34S
148
Chapter 1 Geothermal Systems Geysers:
002 HCO 3 whole rock calcite
a'~c (%0) -20
-10
I
I
0
10
I
Salton Sea/Cerro Prieto: Carbonate host altered carb. host carbonates
Broadlands/Wairakei: calcite
Sediment-Hosted Cortez: carbonate host altered carb. host calcite Carlin: carbonate host altered carb. host calcite
g !
Volcanic-Hosted Pueblo Viejo: carbonaceous sed.
Zoned Polymetallic Veins Creede: Sunnyside: Tui: Casapalca:
carbonates carbonates carbonates calcite
Figure 1.108. Distributions of ~13C in epithermaI deposits (Field and Fifarek, 1985).
values of sulfide sulfur from Green tuff-type epithermal Au-Ag deposits are higher than those from Non-Green tuff-type (Shikazono, 1999b) (Fig. 1.110). The other geochemical and mineralogical features (313C and 3180 of carbonates, salinity of inclusion fluids, Ag/Au total production ratio, association of metals, gangue minerals) are also different in Green tuff-type and Non-Green tuff-type of epithermal Au-Ag deposits (Table !.15). These differences can be explained by the influences of country rocks. The Green tuff-type deposits are affected by marine rocks containing marine sulfates and carbonates and interstitial water of seawater origin with high chloride concentration, resulting in higher ~34S, ~13C, salinity, and Ag/Au total production ratio and enrichment of base metals (Cu, Pb, Zn, Mn, Fe), whereas the Non-Green tuff-type deposits are by subaerial rocks. Sedimentary sulfur and organic carbon were involved in ore fluids responsible for the Non-Green tuff-type deposits, resulting in relatively low ~348 of sulfides and 3 I3C of carbonates. 334S values of epithermal base metal deposits are higher than those of the epithermal Au-Ag deposits and range mostly from +3%o to +7%o (Fig. 1.111 ). Although most of 334S values for base-metal deposits lie in this range, ~34S of composite sample of sulfides from the Motokura Cu-Pb-Zn deposits, Ohmori Cu-Ag deposits, Hosokura Pb-Zn deposits, Sasayama Cu-Pb-Zn deposits and Imai-Ishizaki Cu-Pb-Zn deposits are low, that is, +0.1, +1.8, +2.2, -0.9 and -2.1%o, respectively (Shikazono, 1987b; Shikazono and Shimizu, 1993).
Miocene-Pliocene Hydrothermal Ore Deposits
149 [ ] Ginguro Type
if)
Base metal rich
Ct~
t~ e-
[] Type Disseminated [] Type
~[~
<
"6 E z 1
[~ -7
IoIX~lxlX'gqN]X~ I I ~ I,[ 1,1 I,I 1,
N] -6
.5~.~?~NNN -5
-4
-3
-2
-1
0
1
2
3
4
5
6
-l
8~s(%o)
Figure 1.109. Sulfur isotopic compositions of Neogene Au-Ag vein-type and disseminated-type deposits. Sulfur isotopic compositions on the samples from the Yatani deposits (Sample No. YT26 from Zn-Pb vein ~34S = +3.3%0), and HS72050305-YT1, YT24 and NS-3 from Au-Ag vein (average ~34S = -t-3.3%o)) by Shikazono and Shimazaki (1985) are also plotted. "Base-metal rich" implies the sample containing abundant sulfide minerals but no Ag-Au minerals from base-metal rich deposits and also from Ginguro-type deposits (Shikazono, 1987b).
TABLE 1.15 Some characteristic features (~34S, ~I3c, Ag/Au total production ratio, metals produced, gangue mineraIs) of epithermal Au-Ag vein-type deposits of the Green tuff-type and the Non-Green tuff-type in Japan (after Shikazono, 1996)
334S 3~3C Au produced (metric tons) Ag produced (metric tons) Ag/Au (in wt. ratio) Amount of sulfides (Cu, Pb, Zn)
Green tuff-type
Non-Green tuff-type
high (-1%o to +6%o) high (-7%o to 0%o) 136 2586 high (average 19.1) large
low (-7%o to +2%0) low (-12%o to +2%0) 420 1506 low (average 10.7) small
common common common present
abundant rare rare absent
Gangue minerals
Calcite Rhodochrosite Manganoan calcite Barite
These data indicate that (1) ~348 values of sulfides are different in different mine districts; in the region where thick Green tuff occurs (e.g., Hosokura Pb-Zn deposit, ore deposits in Southwest Hokkaido) some values are relatively high, and ~34S values of the ore deposits in the Non-Green tuff region (e.g., Northeast Hokkaido) (Motokura, Khonomai, etc.) are low, (2) ~34S values of sulfides from small base-metal vein-type deposits (e.g., Imai-Ishizaki, Sasayama) are low, but ~34S values of relatively large base-metal deposits (Taishu, Toyoha, Ani, Osarizawa, Ohe, Jokoku) are high. These data suggest that ~34S values of large deposits are affected by reservoir sulfur in deep part, but sulfide sulfur of small ore deposits is influenced by the surrounding rocks having low ~34S values (basement sedimentary rocks).
Chapter 1
150 15
Non-Green tuff-type
>,1o.
0 c-
OLL
5"
0 q 15" Green tuff-type
o>,10c-
O"
E
LI.. 5-
-15
-5
-10
0
5
10
53'S (permil) Figure 1.110. Frequency of ~348 values of sulfides from the Green tuff-type and the Non-Green tuff-type deposits (Shikazono, 1999b).
Cu.Pb.Zn Vein-Type Deposits ¢--
<
"6 $ E Z i
-8
,
-7
a
i
-6
-5
I
-4
t
i
-3 -2
-1
0
2
3
4
5
6
7 8 8~s (%o)
Figure 1.111. Sulfur isotopic values tbr sulfides from base-metal vein-type deposits (Shikazono, 1987b).
There is another explanation for the variations in 334S values of sulfide sulfur. It was cited that oxidation state (Jb2) and pH of ore fluids are important factor controlling ~34S values of ore fluids (e.g., Kajiwara, 1971). According to the sulfur isotopic equilibrium model (Kajiwara, 1971; Ohmoto, 1972), ~34S of sulfides in predominance
Miocene-Pliocene Hydrothermal Ore Deposits
]
151
Pyrrhotite-bearing
-2-1
0 1 2 3 4 5 6 7
8910
~34s (%°)
Hematite-bearing
•
-2-1
j
•
•
•
.
,
,
,
,
0 1 2 3 4 5 6 7
F]
89'16 8~s (%0)
Figure 1.112. Sulfur isotopic compositionof pyrrhotite-bearing(solid) and hematite-bearing (open) samples from base-metal-richdeposits in Greentuffregion (Shikazono, 1987b).
region of oxidized sulfur species is lower than that in predominance region of reduced sulfur species. Figure 1.112 shows that ~34S of sulfides containing pyrrhotite which formed in the predominance region of reduced sulfur species is lower than that containing hematite which formed in predominance region of oxidized sulfur species; ~348 values for pyrrhotite-bearing samples are less than +4.5%o, while those for hematite-bearing samples are more than +4.5 %o (Fig. 1.112). This is not consistent with equilibrium model (Kajiwara, 1971; Ohmoto, 1972). The ~34Svariation in individual deposits such as Yatani (Shikazono and Shimazaki, 1985), Hosokura, Toyoha (Kiyosu, 1977a; Hamada and Imai, 2000), Ohe (Kojima and Sugaki, 1989) and Taishu (Kiyosu, 1977b) have been studied. These data indicate that ~348 of sulfides do not vary widely in individual deposits. Figure 1.113 shows the distributions of ~34S in epithermal deposits in other countries summarized by Field and Fifarek (1985). ~34S of sulfides from volcanic-hosted epithermal deposits and zoned polymetallic veins are mostly in a range of -5%o to +5%o, typically 0%o (Creede, Sunnyside and Cotqui). This range is similar to that of Japanese epithermal Au-Ag deposits, but lower than that of base-metal deposits. Different interpretations of origin of sulfide sulfur with 0%o have been proposed. For example, Casadevall and Ohmoto (1977) suggest that the sulfur of sulfides in Sunnyside deposits (U.S.A.) originated from evaporite-bearing sedimentary rocks based on the assumption of equilibrium fractionation between H2S and SO ] . in the dominant region of oxidized sulfur species. Kamilli and Ohmoto (1977) also prefer the sedimentary sulfate source at Colqui (Peru), but suggest that an igneous origin is possible. For the Creede deposits, sulfate-sulfide sulfur isotopic equilibrium is not attained and igneous origin seems more likely. Distributions of major epithermal Au-Ag vein-type deposits are shown in Fig. 1.114. Green-tuff-type deposits are defined as the deposits occurring in the Green tuff region and Non-Green tuff type as those occurring in the regions other than the Green tuff region.
Chapter I
152 (~34s
Geothermal Systems Iceland:
magmaticsulfur SO4 sulfates H2S pyrite
-10 I
0
(%0) 10
I
-20 I ! m
Yellowstone: SO4
sulfur H2S Salton Sea: sulfides BroadIands: sulfides Wairakel: SO4 sulfates H2S sulfides
B
B m
I m
Sediment-Hosted Cortez: Carlin:
barite diagenetic py sulfides barite diagenetic py sulfides
32
Volcanic-Hosted Tolfa:
sulfates sulfides Pueblo Viejo: sulfates sulfur sulfides Goldfield: alunite pyrite Z o n e d Polymetallic Veins San Juan Moutains: Creede: barite sulfides Sunnyside: sulfates sulfides Rico: sulfides Ouray: sulfides Tui: barite sulfides Guanajuato: sulfides in country rock sulfides in volc. sulfides in ore Casapalca: sulfides Finlandia: barite sulfides Western Cascades: sulfides Golden Sunlight: barite sulfides
m i
i
n B
45
_.B
m
. . . . . . . . . .
Figure 1.113. Distributions of ~34S in epithermal deposits (Field and Fifarek, 1985).
Figure 1.114 demonstrates that the deposits in northeastern Hokkaido, central Honshu, Sado Island and Kyushu deposits are Non-Green tuff type and those in southwestern Hokkaido, Northeast Japan (Tohoku), the Izu Peninsula, and San-in are Green tuff-type.
Miocene-Pliocene Hydrothermal Ore Deposits 133 °
o,
153
138 °
200 ~
,
%
Hokka~
~
40 ° ~
SW Honshu
35°
~
]
--
,¢ hu '
~
/ ~lzu peninsula
e tuff region t I (submarine volcanic region) ~ Non Green tuffregion I.. • 4 (subaerial volcanic region) Figure 1.114. Distributions of major Green tuff-type (solid circle) and the Non-Green tuff-type (open circle) Au-Ag vein-typedeposits in Japan (Shikazono, 1999b). ~ , , f " / ) MY Kyushu~ /
The Green tuff-type deposits are sometimes hosted not only by the Green tuff formation but also by the rocks overlying the Green tuff formation, implying they are younger than the Green tuff age (ca. 61-15 Ma). Examples of these deposits (Seigoshi, Yugashima, and Toi) occur in the Izu Peninsula, central Honshu. Base metal-rich deposits including epithermal vein-type and Kuroko deposits occur in the Green tuff, while base metal-rich deposits are few in the Non-Green tuff region. Sedimentary rocks often occur as host rocks, footwall rocks and basement rocks in the Non-Green tuff mine area. For example, in southern Kyushu, the Shimanto Supergroup shale is dominant as basement and a host rock for epithermal A u - A g vein-type deposits (e.g., Hishikari). ~34S values of sulfide sulfur from epithermal A u - A g vein type deposits obtained are summarized in Table 1.16 and Fig. 1.109. ~34S value of sedimentary rocks of the Shimanto Supergroup which hosts the Hishikari deposits, southern Kyushu is -12%o (Ishihara et al., 1986) which is considerably lower than those of the Hishikari deposits (+1%o to +2%0) (Shikazono, unpublished. This suggests that sedimentary sulfide sulfur is one of the sources of the sulfides and probably igneous sulfide sulfur is the dominant source. Morishita (1993) showed based on carbon isotopic composition of carbonates that carbon of carbonates in the gold-bearing quartzvein in southern Kyushu was derived from the Shimanto Supergroup shale. Imai et al. (1998) considered that hydrogen in the ore fluids was derived from the Shimanto Supergroup shale based on 3D (--60%0 to -100%o) of inclusion fluids in quartz and adularia of the Hishikari veins. These isotopic
Chapter 1
154 TABLE 1.16
~348 values of epithermal Au-Ag deposits in Japan (Shikazono, 1999b) Green-tuff-type (average +2.9%~, n = 40)
Non-Green tuff-type (average -1.4%~, n = 49) Mine
~348 (%0)
Metals produced
Kohnomai Ryuo Koudou SW H o k k a i d o Katsuyama
-3.5 -7.1 -0.7 -0.9 -7.0 -2.3 -7.2 -5.8
Au, Ag Au, Ag Au, Ag
- 1.3
Au, Ag Au, Ag
-2.8
Au
Sado
Kasuga Hishikari
Iwato
Sappro Eniwa Todoroki
-0.2 -0.2 -14.6 -7.2 -4.2
Au, Au, Au, Au, Au,
Chitose Teine Date Mutsu Innai
Ag Ag Ag Ag, Cu Ag, As
-1.5 -0.9 +1.6 +0.7
Au, Ag
-3.6 +1.0 -7.4 +0.2 +3.4 +5.6 +5.0 -0.5 +1.6 -2.0 --2.8, -1.5 -0.8, +5.5, +3.2, -0.2, -0. l, -2.3 +1.8, +0.4, +0.6, +1.5 +0.3, +0.2, 0.0, -1.1 -0.5, -0.8, -3.7, -5.4
Au Au Au Au Au Au Au Au Au Au
Au, Ag, (Cu)
Handa Yatani
Takatama
+5.1 +4.5 +3.2 +3.2 +3.0 +2.5 +1.8 +4.3 +5.9
Au, Ag
+4.4 +5.2 +4.4 +2.1 +2.9 +3.5 +4.5 -0.4
Au, Ag
Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag, Cu Au, Ag
Ag Ag Ag Ag Ag Ag Ag Ag Ag Ag
Au, Ag
Au, Ag Au, Ag
Au, Ag
Central Honshu
Ashiyasu
+ 1.9
Au, Ag, Cu
+2.9 +2.0 +3.6 +4.9 +3.8 +2.4 +3.2 +5.8 +4.3 +3.2 +3.2 +3.1 +3.1 + 1.0 +0.4 +0.0
Au, Ag
Izu Peninsula
Ohito
Kyushu
Bajo Asahi Hoshino Taio Ohkuchi Yamagano Yamada Arakawa Kushikino Akeshi
Metals produced
NE Honshu
Sado Island
Takachi
Shizukari
Au, Ag
Central Honshu
Nebazawa Haguro Koei Masutomi Suzukura
~34S (%o)
NE Hokkaido
NE H o k k a i d o
Kami-oumu Hokuryu Sanru
Mine
Yugashima Nawaji Jyoren Omatsu Rendaiji Kawazu Okuyama Suzaki Mochikoshi Seigoshi Toi
Au, Ag Au, Ag Au, Ag Au, Ag Au, Ag, Cu, Mn Au, Ag Cu Au, Ag, Cu Au, Ag Au, Ag Au, Ag Au, Ag
+0.1
-0. I Au, Ag
S W Honshu (San- in)
Takeno
+2.4
Au, Ag
Miocene-Pliocene Hydrothermal Ore Deposits
155
data support the view that a part of sulfide ore sulfur originated from sedimentary sulfur in the Shimanto Supergroup shale, although sedimentary sulfide is probably not a dominant source, and large amounts of igneous sulfide sulfur were involved. 3180 values of the Hishikari ore fluids estimated from 3180 values of quartz and adularia and homogenization temperature of fluid inclusions are -6%o to 0%0 (Shikazono and Nagayama, 1993) which are higher than those for the other Japanese epithermal A u Ag ore fluids (Fig. 1.103). These data are interpreted in terms of maturity of hydrothermal system (Shikazono and Nagayama, 1993), boiling of ore fluids (Nagayama, 1993b), a contribution of magmatic water (Matsuhisa and Aoki, 1994), or hydrothermal fluids whose oxygen isotopes exchanged for those of surrounding rocks at a low water/rock ratio (Shikazono and Nagayama, 1993), and a contribution of sedimentary heavy oxygen (Imai and Uto, 2001). Combined isotopic data (3180, 3D, 313C, 334S) indicate that the interaction of hydrothermal solution with the Shimanto Supergroup shale is likely cause for the higher 3180 of the Hishikari epithermal A u - A g ore fluids, compared with those of the ore fluids at main stage of A u - A g mineralization for the other epithermal ore deposits. Shikazono (1988a) has found that ~180 values of late-stage calcite in epithermal ore fluids are higher (0%0 to +10%o), compared with fluid inclusion data (Hattori and Sakai, 1979). Shikazono (1988a, 1989) considered that boiling of ore fluids is important for causing higher 3180 of late-stage calcite. However, it is also likely that this high value seems due to the interaction of hydrothermal solution with sedimentary rocks with high 3~80 values. 334S values of the Green tuff-type are higher than those of the Non-Green tuff-type and are close to ~34S of the Green tuff region (Kuroko and skarn-type deposits) fall in a narrow range of -t-2%o to +7%o and the sulfide sulfur in these ore deposits is called as the "Green tuff sulfur" by Shimazaki (1985). It is inferred by him that the "Green tuff sulfur" was derived from a deep and homogenous reservoir such as magma. However, two sources (igneous source and seawater sulfate) of the "Green tuff sulfur" are also likely because 334S values of base metal deposits (epithermal base metal vein-type and Kuroko deposits) lie between those of igneous sulfur (0%0) and seawater sulfate (+20%o). Kawahata and Shikazono (1988) calculated 334S of hydrogen sulfide (H2S) in ore fluids for midoceanic ridge deposits as a function of seawater/basalt ratio at constant temperature by giving 334S and sulfur content of original basalt. Following their calculation method, we could reasonably explain 3348 of base metal deposits in the Green tuff region (Shikazono, 1987b. However, 334S of the Green tuff-type A u - A g veins are slightly lower than those of the "Green tuff sulfur", indicating that sedimentary sulfur was involved, and degree of contribution of igneous sulfide sulfur was not large. Shikazono (1987b) pointed out a possibility of igneous origin for sulfide sulfur of epithermal A u - A g ore deposits based on average 3348 values of sulfides from these deposits ( - 1%o to +3%0). Smaller involvement of seawater sulfate into epithermal A u - A g ore fluids for the Green tuff-type than the ore fluids for base metal vein-type and Kuroko deposits in the Green tuff region is possible. The mixing of the "Green tuff sulfur" and sedimentary sulfur is likely cause for the difference in 334S of sulfide sulfur of two types of deposits, considering stable isotopic and geologic characteristics of the Green tuff-type and Non-Green tuff-type epithermal A u - A g veins.
156
Chapter 1
25 du N
20, 0
"~ 15. i •
~, 10-
0 a D o
5-
•i
~
•
Im~ •
•
•
mOj mmn--~ m •
0 o
-15
-10
Non-Grenn tuff type 0 Sado A Fuke Hishikari
.~
613C of Carbonates (%~)
b
Green tuff type • Nawagi • Kawaza-Kakehashi • Yugashima
• •
Yatani Seigoshi
Figure 1.115.3tSO-~t3Cof carbonates fromepithermal vein-typedeposits in Japan (Shikazono, 1999b).
313C values from both types of deposits are different (Fig. 1.115). 313C of the Non-Green tuff-type carbonates are relatively low, suggesting a contribution of organic carbon in sedimentary rocks in the mine areas. Although the ~13C values of the Green tuff-type vary widely, the data lie between igneous value (-7%0 to -8%0) and seawater value (-1%o to -4%o), suggesting a contribution of these two sources. Total metal (Au, Ag, Cu, Pb, Zn, Mn) production during the past, and Ag/Au total production ratio of major epithermal Au-Ag deposits (Green tuff-type and Non-Green tuff-type) are summarized in Table 1.17. Large amounts of Au and Ag have been produced from the Non-Green tufftype (e.g., Hishikari, Sado, Kohnomai, Kushikino). Ag/Au total production ratio of the Non-Green tuff-type (average 10.7) is lower than for the Green tuff-type (average 19.1). These differences can be explained by the HSAB principle by Pearson (1963, 1968). This principle indicates that HS- and H2S are likely to form complexes with the metals enriched in the Non-Green tuff-type (Au, Hg), whereas CI prefers to form complexes with the metals concentrated in the Green tuff-type (Ag, Pb, Mn, Fe, Cu). Occurrence of gangue minerals in both types of deposits is different. For example, Mn minerals (Mn carbonates, Mn silicates) occur abundantly in the Rendaiji, Yugashima, Yatani, and Todoroki epithermal Au-Ag vein-type deposits in the Green tuff region but not in the Non-Green tuff-type. 334S values of barite from these deposits are high (+18%~
Miocene-Pliocene Hydrothermal Ore Deposits
157
TABLE 1.17 Tonnages of gold, silver and the other associated metal, silver/gold ratio, K-Ar ages and host rocks for the Te-type and Se-typc epithermal gold deposits (Shikazono et ah, 1990) Deposit
Au (ton)
Ag (ton)
Ag/Au
Sanru
6.7
40
6.0
Koryu Chitose Yatani
0.76 22.8 1.7
22.2 105 64
29.4 4.7 39
Takadama Nebazawa Seigoshi Omidani Sakoshi-Odomari Kushikino Hishikari
28.8 1.0 13.5 0.3 1.1 55.2 21.7
279.9 65 455 79 9.7 456.2 14.3
Date Kobetsuzawa Teine Chitose Mutsu Osorezan Sado
<1 <1 10.4 22.8 <1
Other metals
Year of production (ton)
K-Ar ages
Host and country rocks
1925-1974
12.4 4- 0.6
1903-1957 I93(>-1974 1870-1974
1.0 4- 0.3 4.7 3.34.0.3
Rhyolitic tuff, tuff breccia, shale Shale Propylite Acidic tuff, shale
9.7 65 34 267 8.8 8.3 0.6
1429-1974 1942-1974 1935-1976 1914-1974 1977-1982 1965-1986 1983-1990
8.4 5.0-5.7 1-3.7 66-68
<1 <1 62.6 105 <1
0.7-2 48.1 6.0 4.7 0.1
1932-1975 1955 1932-1971 1936-1974 1940
5.2:t:0.4
82.9
2404.3
33
Cu: 5400
1601-1988
Kawazu
5.4
272.8
1-5
1915-1959
Okuyama Suzaki Chugu Takeno
<1 2.0 <1 4.6
<1
Cu: 1054 Mn: 15840 Cu: 1250
13.44-0.514.5±0.5 22.1 4-0.7 24.4 4- 0.8 1.44-0.31.5 4- 0.3
<1 90.0
Agawa Kato
<1 <1
<1 <3
lriki Yamada Fuke Okuchi Kushikino
<1 <1 2.4 21.2 55.2
<1 <7 1.5 I5.7 456.2
Se-type
Cu: 70 Cu: 1270 Pb: 29210 Zn: 58420
4.0 4- 0.3 0.97 4- 0.041.54.0.3
Tuff, shale Rhyolite Andesite, diorite porphyry Slate, sandstone Dacitic tuff Andes±re, propylite Shale, andesite
Te-type
Cu: 7560.8 Cu: 70
0.1 20
10-I2 I054 0.7 0.7 8.3
Cu: 200 Pb: 500 Zn: 100 Cu: 1.3
1912-1963 1914-1941 1938 1868-1949
4.7
17.9 4- 4.518.24-4.6
1933-1945 1933-1954
1933-1937 1916-1955 1896-1947 1905-1974 1865-I 986
Tuff Propylite Propylite Propylite Liparitic tuff Dacite Dacite, rhyolite shale, dacitic tuff
Propylite, rhyolite Propylite, dacite, basalt Propylite, rhyolite Tuff Granite, tuff, andesite Rhyolite, dacitic tuff Porphyrite
0.453 -t-0.018 1.4 4- 0.2 1.i :t- 0.5 4.0 :k:0.3
Tuft, andesite, liparite Propylite Propylite Andesite, rhyolite Propylite, andesite
158
Chapter 1
to +29%o) (Watanabe and Sakai, 1983; Shikazono et al., t990), suggesting that sulfate sulfur of barite was derived from sulfate in the submarine volcanic and pyroclastic rocks (gypsum, anhydrite, sulfate ion in interstitial water) (Watanabe and Sakai, 1983). Carbonates are found both in the Green tuff-type and the Non-Green tuff-type, but large amounts of carbonates (dominantly calcite) occur in the Non-Green tuff-type such as the Kushikino and Kohnomai epithermal Au-Ag vein-type deposits. ~13C and 8180 of carbonates from epithermal vein-type carbonates are plotted in Fig. 1.106. 313C values from both types of deposits are different, gl3C values of the Non-Green tuff-type carbonates are relatively low, suggesting a contribution of organic carbon in sedimentary rocks in the mine areas. Although the ~13Cvalues of the Green tuff-type vary widely, the data lie between igneous value (-7%o to -8%0) and seawater value (-1%o to +4%o), suggesting a contribution of these two sources.
834S and 8180 of sulfates. Watanabe and Sakai (1983) have analyzed 3348and 3180 of barite, anhydrite and gypsum in the epithermal vein-type deposits. They found that (1) the hydrothermal vein sulfates are characterized by more scattered distributions of 334S values ranging mostly from -t-10%o to .1.20%o and 8180 mostly from 0%0 to +14%o than those of Kuroko sulfates (section 1.3.3) (Fig. 1.43). In Fig. 1.43, 334S and 8180 of sulfates from epithermal vein-type deposits (Watanabe and Sakai, 1983) are plotted. These data show that 334S (mostly from .1.24%o to -1-t-37.8%o)and 3180 of barite (0.1%o to ,1,1,18.7%o)from epithermal Au-Ag-Te vein-type deposits are higher than that of epithermal base-metal vein-type deposits (~34S; .1.16.0%o to +24.6%o, 3180; +2.1%o to +12.1%o). These data could be explained by: the sulfur of barite from epithermal Au-Ag-Te deposits came both from volcanic gas (SO2) and marine sulfate, but that of epithermal base-metal deposits came from marine sulfate and oxidation of H2S. 1.4.3.6. Lead isotopes Lead isotopic data on the epithermal deposits together with Kuroko deposits are plotted in Fig. 1.116 (Sato and Sasaki, 1973; Sato et al., 1973, 1981; Sato, 1975; Sasaki et al., 1982; Sasaki, 1987; Fehn et al., 1983). It is evident that lead isotopic compositions of epithermal vein ores are more scattered than Kuroko ores, although averaged values are similar to the Kuroko ores. This variation seems to be due to the difference in crustal materials underlying the ore deposits; Lead isotopic compositions of different ore deposits which formed at different ages in the same district show the same values (Sasaki, 1974). 1.4.3.7. Rare earth elements (REE) Kato et al. (1990) analyzed carbonates from epithermal Cu, Zn-Pb and Au-Ag vein-type deposits in Japan for rare earth elements (REE) and found that (1) calcite from Cu-type is characterized by high La/Yb, high REE concentration and negative Ce anomaly is small, and (2) calcite from Zn-Pb type is characterized by high La/Yb, low REE and negative Ce anomaly. These data suggest that ore fluids for Cu-type were generated under low water/rock ratio condition or were influenced by magmatic water, while meteoric water component was dominant for the Zn-Pb type.
Miocene-Pliocene Hydrothermal Ore Deposits 39.0
,
,
%
159 ' !
i
: ............. l "CO
io
C, .04
uo
:d
' o i: t°
°"~
:
38.5 L .............
.t
SRM-981 ,. ,tie
[.r,
206/204
t*"
15.7
!
i
!
i
/k.
!
!
!
I
o TA
'O
0
15.6 • o e4 SRM-981
F~"7"';q :oL:z.?.°?...}
n W
H
.d"
15.5
206/204 I
15./,
16.9
,
I
17.0
~
t
18.3
I
18./,
,
I
l&5
I.
187
Figure 1.116. Lead isotopic variation in Japanese Neogene ores. The majority of data fall in a relatively narrow range which is no more than twice the experimental uncertainty indicated by the replicate analyses of NBS-SRM-981 standard (Sasaki et aI., 1982).
The REE characterics of calcite from the A u - A g type are variable. For example, calcites from Sado A u - A g vein, one of the largest A u - A g deposits in Japan have both signatures of meteric water and magmatic (or igneous) contributions. Positive Eu anomaly is only found in calcite containing low REE from A u - A g type (Seigoshi deposit) (Shikazono, unpublished). 1.4.4. Se- a n d T e - t y p e A u - A g d e p o s i t s As shown in Fig. 1.117, Se-type and Te-type epithermal A u - A g vein-type deposits are located in the Cretaceous-Quaternary volcanic terrane of Japan (e.g., northeast and southwest Hokkaido, middle Honshu, south Kyushu). Some Te-type deposits are located in regions similar to the Se-type deposits. Sometimes, Te mineralization is associated with the Se-type deposits, though Te minerals usually do not coexist with Se minerals. However, rarely, Te minerals coexist with Se minerals in the Te-type deposits (e.g., Teine, Suzaki, Kawazu, Iriki) on a polished section scale. For example, coexistence of native Te and Se-bearing tetrahedrite is found at Teine. Generally, Te mineralization occurs at
Chapter 1
160
Teine
I oh,os ( 200km
-Kobe~uzawa Date
Osorezan
Mutsu ."
4o~u
Sado
Chugu
a :no
Ornidani
0 0 Kawazu
1 ~=1
Solgoshi Iriki
okuchl MTL
I
~ . 133°
2
I
138°
~
4
s ['7"1
3 r!'N 615 1
Hlshikari Yamanda Figure 1.117. Map showing the distribution of Se-type and Te-type epithermal gold deposits in Japan. 1: Green-tuff and subaerial volcanic region of Tertiary/Quaternary ages, 2: Main Paleozoic/Mesozoic terrane, 3: Main metamorphic terrane, 4: Te-type deposits, 5: Se-type deposits, 6: Te- and Se-bearing deposits, ISTL: Itoigawa-Shizuoka tectonic line, TTL: Tanakura tectonic line, MTL: Median tectonic line (Sbikazono et al., 1990). Kushikino
~
a higher level than Se mineralization in the same mine district. Aoki (1988) reported that gold precipitation is currently taking place from the Osorezan hot springs, north Honshu, Japan. At Osorezan A u - T e minerals (e.g., krennerite, coloradoite) are found at very shallow levels from the surface but no Se minerals have yet been identified (M. Aoki, personal communication, ] 988). K - A r ages data on adularia and sericite in the veins and altered host rocks indicate that ages of mineralization vary widely, ranging from 1 Ma to 68 Ma and from 1 Ma to 24 Ma for the Se-type and Te-type, respectively (Tables 1.17 and 1.18). The total production of gold, silver and other associated base metals and silver/ gold production ratios from these deposits are summarized in Table 1.17. In addition to gold and silver, lead, zinc and manganese have been produced from some of the Se-type (e.g., Yatani) and copper has been produced from some of the Te-type (e.g., Teine, Kawazu). Total tonnage of production of Au and Ag from the Se-type is greater than
Miocene-Pliocene Hydrothermal Ore Deposits
161
TABLE 1.i8 Summary of geologic, mineralogic and geochemical characteristics of the Te-type and Se-type deposits (Shikazono et al., 1990) Se-type Associated metals Ag/Au Host rocks
Au, Ag, Pb, Zn, Mn More than l0 Sedimentary rocks (dominantly shale), volcanic rocks (dacite and andesite) Form of deposits Vein Age of mineralization Late Cretaceous-Quaternary Opaque minerals Argentite,polybasite, naumannite, aguilarite, Ag-rich electrum, pyrargylite, sphalerite (FeS; 1-5 wt. %), galena, tetrahedrite, chalcopyrite, pyrite Gangue minerals Quartz (fine-large grained), adularia, illite/smectite, chlorite/smectite, calcite, rhodochrosite Homogenization 150-270°C temperatures ~34S of sulfides - 8 to +5% ~348 of sulfates Sulfur activity low Oxygen activity low pH high
Te-type Au, Ag, Cu, Bi, (Hg), (TI) Less than I0 Volcanic rocks (dacite and andesite) Vein, massive Miocene-Present Hessite, petzite, native Te, sylvanite, Au-rich electrum, chaIcopyrite, pyrite, marcasite, tetrahedrite, enargite, bismuthinite, sphalerite (FeS: 1 wt. %), galena Quartz (very fine-grained), barite, illite, kaolinite, adularia 200-300°C - 3 to +7%0 +18 to +29%o intermediate intermediate intermediate
those of the Te-type. A g / A u production ratio (in weight) from the Te-type and Se-type vary widely, generally less than 10 for the Te-type and more than 10 for the Se-type. Host rocks for the Se-type and Te-type epithermal A u - A g deposits are summarized in Table 1.17. In general, the dominant host rocks for both deposits types are intermediate and felsic volcanic rocks. Sedimentary rocks (usually shale) sometimes host the Se-type (e.g., Sanru, Kohryu, Takadama, Ohmidani, Hishikari), but never host the Te-type deposits. All o f the Se-type are vein in form, and most of the Te-type are also vein in form (Table 1.18). However, some Te-type (e.g., Kobetsuzawa, Date, Suzaki, Iriki) are stratabound or massive. Total production of Au from the Te deposits is small. Size (length and width of the vein) of the Se-type vein is generally greater than the Te-type vein. Temperatures of formation can be estimated from homogenization temperatures of fluid inclusion. The typical range o f temperature o f formation for the Se-type and Te-type is 150-270°C and 200-300°C, respectively (Tables 1.18 and 1.19). The pressure correction to obtain true formation temperatures based on homogenization temperatures of fluid inclusions has not been carried out because evidence for boiling is found in the fluid inclusions from some deposits such as Chitose (Yajima, 1979), Nebazawa (Enjoji and Nakayama, 1982); furthermore, homogenization temperatures of fluid inclusions are in good agreement with electrum-sphalerite temperatures (Shikazono, 1985d), indicating that the pressure o f formation is close to the vapor saturation curve. Electrum-sphalerite geothermometer can be used to estimate the temperature of formation for some deposits
Chapter 1
162
TABLE 1.19 Temperatures of formations of the Te-type and Se-type epithermal gold deposits estimated from fluid inclusion homogenization temperatures (Shikazono et al., 1990) Deposit
Homogenization (°C) Range
Average
No. of data
Sanru Koryu Chitose Yatani Takadama Nebazawa Seigoshi Omidani Hishikari Kushikino
200-300 190-340 145-349 209-273 162-240 215-260 !95-243 148-178 90-260 210-250
250 255
419
241 200 250 225 160 195 235
30 100 3600 10 232 100
Te-type deposits Teine Chitose Sado Kawazu Okuyama Takeno Kato Fuke Okuchi Kushikino
180-240 220-260 247-305 206-281 193-284 210-270 190-330 220-280 164-240 260-310
265 236 261 250 250 245 230 280
34 31 53 110 16
Se-type deposits
17
of the Se-type but not for the Te-type, because the assemblage of electrum-argentite (or acanthite)-sphalerite-pyrite is absent in the Te-type (Shikazono et al., 1990). The dominant opaque minerals from the Se-type deposits are Se-bearing A g minerals (argentite, acanthite, polybasite, naumannite, aguilalarite, pyrargyrite), electrum, tetrahedrite, sphalerite, Se-bearing galena, chalcopyrite and pyrite (Table 1.20). The dominant opaque minerals from the Te-type are hessite, native Te, petzite, pyrite, marcasite, enargite and bismuthinite (Table 1.21 ). These Te minerals are not present in the Se-type deposits. Very rarely Se-bearing minerals (tetrahedrite, bismuthinite) occur in the Te-type deposits. Generally, sulfide minerals except for pyrite and marcasite are very poor in amount in the Te-type. On the other hand, sulfide minerals such as argentite, sphalerite and galena are abundant in the Se-type. The chemical compositions of opaque minerals (sphalerite, electrum) are different in two types o f deposits. The FeS content of sphalerite from vein deposits of the Te-type is generally lower than that of the Se-type (Fig. 1.118). However, FeS content of sphalerite from massive deposits of the Te-type (Kobetsuzawa, Suzaki) is high, ranging from 1 to 7 mol%. The Ag content o f electrum from the Se-type is higher than that from the Te-type (Fig. 1.119).
TABLE t.20 Dominant opaque and gangue minerals from the Se-type epitb_ermal gold deposits (Shikazono et al., 1990)
¢5
Deposit
Se-bearing minerals
Opaque minerals
Gangue minerals
Sanru
aguilarite, naumannite, poiybasite, pyrargyrite, stephanite
electrum, miargyrite, chalcopyrite, lahore, arsenopyrite, marcasite, pyrite, sphalerite, stibnite cinnabar
quartz, adalaria, kaolinite, seficite, calcite
Koryu
aguilarite, pearceite, polygasite, proustite, pyrargyrite
electrum, miargyrite, native silver, chalcopyrite, lahore, hematite, magnetite, pyrite, galena, sphalerite
quartz, adularia, johannsenite, chlorite, kaolinite, vermiculite, Mn-calcite
Chitose
aguilarite, argentite, pearceite, polybasite, proustite
electrum, chalcnpyrite, lahore, pyrite, galena, sphalerite
quartz, adularite, chlorite, sericite, calcite
Yatani
argentite
electrum, chalcupyrite, marcasite, pyrite, pyrrhotite, galena, sphalerite
quartz, adularia, chlorite, sericite, rhodochrosite
Takadama
aguilarite, freibergitc, naumannite, pearceite, polybasite, pronstite, pyrargyrffc, stephanite
electrum, sternbergite, chalcopyrite, fahore, marcasite, pyrite, galena, sphalerite
quartz, adularita, kaolinite
Nebazawa
argentite, polybasite, proustite, pyrargyrite
electrum, mirargyrite, native silver, chalcopydte, covellite, fahore, arsenopyrite, marcasite, pyrite, galena, sphalerite
quartz, adularia, sericite, calcite
Seigoshi
argentite
electrum, pearceile, polybasite, pyrargyrite, stephanite, chalcopyrite, fahore, galena, sphalerite
quartz, adularia, inesite, xonotlite, chlorite, mixed layer clay mineral, sericite, calcite, rhodochrosite
Omidani
aguilarite, argentite, naamannite
electrum, native silver, pearceite, polybasite, chalcopyrite, lahore, pyrite, galena, sphalerite
quartz, adularia, chlorite, ankerite, calcite
SakoshiOdomari
aguilarite, argentite, jalpaite, naumarmite, polybasite, pyrargyrite
electrum, mckinstryite, native silver, chalcocite, chalct~pyrite, covellite, fahore, famatinite, native copper, arsenopyrite, hematite, pyrite, pyrrhotite, galena, sphalerite
quartz, kaolinite, sericite
Hishikari
naumannite
electrum, chalcopyrite, marcasite, pyrite, galena, sphalerite, stibnite
quartz, adularia, montmorillonite
Kushikino
aguilarite, argentite, naumannite, polybasite, pyrargyrite, stephanite
electrum, pyrostilpnite, chalcocite, ehalcopyrite, covellite, lahore, marcasite, pyrite, galena, sphalerite
quartz, adularia, mixed layer clay mineral, smectite, calcite, zeolite, truscottite
d
r~ ,~
~"
TABLE 1.21
g
Dominant opaque and gangue minerals from the Te-type epithermal gold deposits (Shikazono et aL, 1990) Deposit
Te-bearing minerals
Date
calaverite, coloradoite, electrum, lahore, pyrite tellurantimony altaite, frohbergite, hessite, native dyscrasite, chalcopyrite, hematite, tellurium, rickardite, sylvanite, marcasite, pyrite, pyrrhotite, tellurantimony galena, sphalerite, realgar, stibnite lhhore, gold-fieldite, hessite, native bismuthinite, lahore, hakite, native electrum, bornite, chalcocite, tellurium, rickardite, stuetzite, tellurium, stuetzite chalcopyrite, emplectite, enargite, sylvanite lahore, luzonite, hematite, marcasite, pyrite, galena, sphalerite, orpiment, realgar, stibnite hessite, petzite electrum, cha]copyrite, lahore, pyrite, galena, sphalerite krennerile, native tellurium chalcopyrite, fahore, arsenopyrite. marcasite, pyrite coloradoite, krennerite marcasite, pyrite, orpiment, realgar hessite electrum, polybasite, pyrite empressite, fahore, gold-fieldite. chalcopyrite, fahore, hemusite, kawazulite, native tellurium, hessite, kawazulite, native stannoidite, hematite, pyrite, poubaite tellurium, poubaite, nckardite, sphalerite stuetzite, sylvanke, tellurobithmuthite, tetradymite altaite, hessite, tetradymite electrumm chalcopyrite, native copper, hematite, pyrite, sphalerite calaverite, hessite, kawazulite, clausthalite, kawazulite, native chalcopyrite, marcasite, pyrite, kostovite, krennerite, native tellurium, poubaite, pyrite sphalerite tellurium, petzite, poubaite. sylvanite hessite electrum, argentite, chalcopyfite, electrum fahore, marcasite, pyrite, galena, sphalerite hessite, petzite
Kobetsuzawa
Teine
Chitose Mutsu Osorezan Sado Kawasu
Okuyama Suzaki
Chugu
Takeno
Se-bearing minerals
Opaque minerals
Gangue minerals quartz, anatase quartz, scricite, zeolite, anatase
quartz, sericite, calcite. rhodochrosite, barite
quartz, adularia, chlorite, sericite, calcite quartz quartz, barite quartz, adularia quartz, sericite, barite, anatase
quartz, chalorite, sericite quartz, barite, anatase
quartz.
quartz, adularia, chlorite, kaolinite, sericite, calcite, dolomite
e~
t"5.
TABLE 1.21 (continued) Dominant opaque and gangue minerals from the Te-type epithermal gold deposits (Shikazono e~ al., 1990) Deposit
Te-bearing minerals
Agawa
calaverite, telluro-bismuthite, tetradymite hessite, sylvanite
Kato lriki
fahore, gold-fieldite, native tellurium
Yamada
hessite, tetradymite
Fuke
hes~ite
Okuchi
hessite
Se-bearing minerals
lahore, famatinite
Opaque minerals
Gangue minerals
electrum, bornite, chalcopyrite, hematite, pyrite, molybdcnite electrum, chalcopyrite, pyrite, galena, sphalerite argentite, chalcopyrite, fabore, famatinim, arsenopyrite, marcasite, pyrite, pyrrhotite electrum, chalcopyrite, pyrite, sphalerite electrum, chalcopyrim, hematite, pyrite, galena, sphalerite electrum, argentite,
quartz, sericite, calcite quartz quartz, kaolinite, anatase
quartz, calcite quartz, adularia, chlorite, kaolinite, smectite quartz, adularia,
Chapter 1
166
In the Se-type gangue minerals comprise quartz, adularia, illite/smectite interstratified mixed layer clay mineral, smectite, calcite, Mn carbonates (manganoan calcite, rhodochrosite), Mn silicates (inesite, johansenite) and Ca silicates (xonotlite, truscottite). In comparison, the Te-type contains fine-grained, chalcedonic quartz, sericite, barite, adularia and chlorite/smectite interstratified mixed layer clay mineral. Carbonates and Mn minerals are very poor in the Te-type and they do not coexist with Te minerals. Carbonates are abundant and barite is absent in the Se-type. Grain size of quartz in the Te-type is very fine, while large quartz crystals are common in the Se-type. Hydrothermal alteration patterns in the Se-type and Te-type deposits are not well defined due to a lack of detailed studies of hydrothermal alteration. However, it is evident that propylitic alteration is widely developed in Se-type gold mine districts (Shikazono, 1988a). For example, chlorite and chlorite/smectite interstratified mixed layer clay mineral are abundant in the country rocks which host the Seigoshi, Fuke, Kushikino and Hishikari deposits (Inome et al., 1981; Shikazono, 1985a; Shikazono et al., 2002). Alteration characterized by the presence of adularia is well developed in some Se-type mine districts (e.g., Seigoshi, Hishikari) (Shikazono, 1985b). Argillic alteration overlies the Kushikino, Seigoshi and Takadama vein deposits (Yagyu, 1954a,b; MITI, 1979). The country rocks in Te-type mine districts suffered argillic and sericite alterations and silicification. For example, such alterations are developed in the Kawazu Te-type gold
90 80
E] Te type deposits (vein) [] Te type deposits (disseminated) [] Se type deposits
70 60
~'50 ~ 4o Ii 30 20 ¸ 10
2
3
4
s
6
7
8
1'0 1112 13t;
is
1617
geS Mol% Figure 1.118. Frequency (number of analyses) of FeS content (mole fraction of FeS) of sphalerite from the Te-type and Se-type deposits (Shikazono et al., i990).
Miocene-Pliocene Hydrothermal Ore Deposits
167
Te Type deposits • n=5 Date Okuyama ~l-~n=10 • n=2 Kushikino Agawa Chugu ,~'H. n=15 Fuke = ~n=69 Okuchi n=10 en=17 Chitose : z = n=6 Takeno == n=l Sado Kato SeType deposits
• n=l ~ n=61
Sakoshi Hishikari Koryu Chitose Sanru Kushikino Takadama Yatani Nebazawa Omidani
n=20~
x =
=
~ n=22
4
=
~ n=8
n=41 = : .~ n=44
I
0
I
I
I
I
I
50
I
I
I N Ag
I
I
100
Figure 1.119. Ag content (atomicfraction of Ag) of electrum from the Te-typeand Se-typedeposits (Shikazono et al., 1990). n: number of analyses.
mine district in Izu Peninsula (Watanabe and Nagai, 1986). Advanced argillic alteration closely related to gold mineralization has not yet been recognized in the Se- and Te-type gold deposits. Sulfur fugacity can be estimated from the homogenization temperatures of fluid inclusions, FeS content of sphalerite coexisting with pyrite, Ag content of electrum coexisting with argentite (or acanthite) and the assemblages of sulfide minerals. Typical ranges of sulfur and oxygen fugacities for the Se-type and Te-type are shown in Fig. 1.120. It is likely that sulfur and oxygen fugacities for the Te-type are higher than those for the Se-type at the same temperature conditions. Probable ranges of oxygen fugacity and pH for the Se-type and Te-type are shown in Fig. 1.121. The pH and oxygen fugacity of ore fluids responsible for the formation of the Se-type appears to be lower and higher, respectively, than for those of ore fluids responsible for the formation of the Te-type (Fig. 1.121). Sulfur isotopic compositions (~34S) of sulfides and sulfate (barite) from the Setype and Te-type are summarized in Fig. 1.122. Almost all ~34S values from the Se-type and Te-type fall in a range from -3%o to +6%o (Fig. 1.122). In general, the ~34S values from the Se-type are similar to those of the Te-type. However, some ~348 values from the Se-type are lower than those from the Te-type. The ~34S values of sulfate (barite) sulfur from the Te-type range from +18%o to +29%0 (Fig. 1.122).
Chapter 1
168 -6
#
!
i
i
t~
COVELLITE DIGENITE
~_\'~
PYRITE+BORNITE
/I
ENARGITE
!
FAMACHtNITE F , \ Te-type ,o.;"~
--
......~
\/
~ /
\
1 mole%FeS in SPHALERITE ['-.t~~',"~'~-~-I" ," • ,it."
0
4%
. _//"
i"" s -tyr, e
7-'- % . . ; % . . d . .14
-38
g:
, -37
-35
-33
i -31
Log ao2 Figure 1.I20. Probable ranges of sulfur and oxygen activities for the Te-type and Se-type deposits. The diagram was constructed mainly based on Heald et al. (1987). Temperature = 250°C (Shikazono et aI., 1990).
The above-mentioned geologic, mineralogic and geochemical characteristics of the Se-type and Te-type are summarized in Table 1.18. These characteristics can be used to reconstruct the structure of the fossil geothermal system responsible for the formation of the Se-type and Te-type epithermal gold mineralizations. Figure 1.123 shows the relative sites of Se-type and Te-type epithermal gold deposition in a fossil geothermal system. The Te-type is considered to form a a site closer to the volcanic centre than the Se-type. SO2 gas derived from magma disproportionates in the presence of water to form sulfate ion and hydrogen ion according to the following reactions (Holland and Malinin, 1979). 4SO2 + 4 H 2 0 --+ 3 H2SO4 q--H2S
(1-45)
H2SO4 --+ 2H + + SO 2-
(1-46)
The sulfate ion generated by these reactions precipitated as barite in some Te-type deposits. This hydrolysis reaction can explain the 334S values of barite and sulfides in the Te-type deposits. Some of the lower 334S values of sulfide sulfur in the Se-type deposits than the Te-type deposits could be due to the incorporation of sedimentary sulfur in the host rocks of the Se-type deposits. However, sulfide sulfur of the igneous source was also involved in the ore fluids responsible for the Se-type. Relatively higher sulfur and oxygen fugacities and lower pH and higher temperature of ore fluids responsible for the Te-type than the Se-type may have been caused by a contribution of magmatic (or igneous) sulfur to the ore fluids responsible for the Te-type. However, it cannot be ruled out that seawater sulfate was involved in the ore fluids responsible for the Te-type, because ~34S of barite
Miocene-Pliocene Hydrothermal Ore Deposits -30
I
W!W
I
I
169
J
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I
l
_
~
~L 7 -~,,~
, "1
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I
--I>zl~
#.,
I
]
-
/t
~N'~'~ .
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- , o
-
.
.
.
s
.¢"
,I
0,1%
E
~,~~t '
.
i 2
4
pH
6
8
1
10
Figure 1.121. Probable ranges of oxygen activity and pH for the Te-type and Se-type deposits. The diagram was constructed mainly based on Barton et al. (1977) and Heald et al. (1987). Temperature = 250°C, ICS = 0.02 molal, Salinity = 1 molaI with N a / K (atomc ratio) = 9. Dotted area: Te-type, Hatched area: Se-type (Shikazono et al., 1990).
are close to seawater sulfate value and Te-type deposits tend to occur in Green Tuff regions characterized by thick marine sedimentary and volcanic rocks. Advanced argillic alteration is not found in either the Te-type or Se-type deposits. Sericitization, silicification and argillic alteration are common in the Te-type mine districts. Adularia is a characteristic mineral of adularia-sericite-type (Heald et al., 1987). Thus, the Te-type appears to be intermediate type between the acid-sulfate-type and adularia-sericite-type. The Se-type belongs to low sulfidation-type, essentially adulariasericite-type. Almost all Te-type deposits belong to low sulfidation-type but some have characteristics of the high sulfidation-type. Acid-sulfate, high sulfidation deposits form just above the volcanic centre. However, the Te-type form further from the site of the volcanic centre, though close to the centre than the Se-type. The Se-type may occur in basement (mainly sedimentary rocks) as well as young volcanic rocks. The models for epithermal system by Buchanan (1981), Giles and Nelson (1983), and Berger (1985) emphasize vertical change from base metal-type deposits at the deeper part to precious
Chapter 1
170
(a)
[ ] Te type deposits [ ] Se type deposits
¢o
R
"6 E Z
-8
8
.I.I "1"1 "1-1 [ • [/Ij1 "1 "1" • I . I , I . I A J~J I - V I'1/1"1"1"1" • I-I "I "Pq [2t , I Z I ~ I - I ,, V l t l - I - . I - I t l - J : l " - I . I . I . I - I , I , . I . I , -6 -5" -4 -3 -2 -1 0 1 2 3 4 5 6 7 •
..Q
-7
(b)
BARITE
"6 E
[]
Teine
]
Kawazu
]
osorezan
27
F~. 28 29 ~'S(%o)
y.
, l'~, 18 19
20
. if'l, 21 22
23
. ITIQ~TI 24 25 26
Figure I. 122. Frequency (numberof analyses) of (a) ~348 of sulfide and (b) sulthte (barite) minerals(Shikazono et al., 1990). metal-type deposits at the shallower part. However, the above discussion emphasizes lateral changes of epithermal ore-types in the geothermal system, as shown in Fig. 1.123. It is worth comparing the fossil geothermal system associated with epithermal gold deposition with active geothermal systems. Henley and Ellis (1983) and Henley (1985) showed that sulfate-type hot springs occur at the volcanic centre and upper level of geothermal system, while bicarbonate hot springs occur on the margins of the geothermal system. Lateral flow of acidic hot waters from upper level of active geothermal system of high relief, with significant lateral flow, appears to be consistent with the distribution of the Te-type and Se-type. Carbonates are common in the Se-type, and barite is found in the Te-type. The occurrence of carbonates and barite suggests that ore fluids for the Se-type and Te-type are bicarbonate-rich and sulfate-rich, respectively. 1.4.5. Depositional mechanism and origin of ore fluids
1.4.5.1. Depositional mechanism It is widely accepted that boiling of ore fluids took place in the epithermal A u Ag mineralization system from the fluid inclusion data (Nakayama and Enjoji, 1985;
Miocene-Pliocene Hydrothermal Ore Deposits
Sotfatara "
Propylltlc / Z
171
""
~'
Figure 1.123. Schematic model for the formations of the Te-type and Se-type epithermal gold depositions in the fossil geothermal system. Reference: Henley and Ellis (1983) (Shikazono et al., 1990).
Shikazono, 1985a) and stable isotope data (Shikazono, 1988a, 1989). Generally, f s 2 and fo2 increase in response to boiling (Drummond, 1981). If boiling occurs, H2 gas escapes from ore fluids faster than other gaseous species, resulting to higher f02 and fs2- Thus, it seems likely that Au-rich electrum tends to precipitate under the high fs2 and f02 conditions from the ore fluids from which substantial amounts of vapor released. However, if boiling of ore fluids having high CO2 and CH4 concentrations occurs, fs2 and f02 do not considerably increase (Drummond, 1981). This type of boiling or gas loss could lead to the deposition of sulfide minerals (sphalerite, galena, argentite etc.) due to a pH increase (Drummond, 1981; Hedenquist and Henley, 1985). This pH increase is not favorable for the deposition of Au when thio-Au species are dominant among dissolved Au species because solubility of Au due to thio complexes increases with the pH increase (Fig. 1.102). If electrum coprecipitates with argentite under low fs2 and f02 conditions, although the amount of electrum precipitated might be small, considering the solubility of electrum under such conditions, electrum tends to contain high Ag content. Not only the deposition of Au, but also the deposition of Ag has to be taken into account in order to consider the depositional mechanism of electrum. Deposition of Ag from the ore fluids in which Ag chloro complexes (e.g., AgCI~-) are dominant Ag species may be controlled by the following reaction. AgCI~- + 1/2 H20 -+ Ag ÷ 2 C1- + H + ÷ 1/4 02
(1-47)
Thus, a pH increase could be favorable for the deposition of Ag. This may imply that the Ag content of electrum is high when substantial amounts of CO2 loss occur. The CO2 concentration of ore fluids responsible for epithermal base-metal vein-type deposits
Chapter 1
172 NAg
o2
1.0
d
"3 0.8
0.6
23 0
•
16 21 15 • j4 19 ,w
•
•
11 e7 • elO o6 120
*17
o8
25 • 22 200 ~8 e24 26
0.4
5 e4
9 13
27° 0.2
30
,28
r'l
-,
29 Q
;
+,
i
~
i log ( A g / A u )
Figure 1.124. Ag/Au total production ratio from each mine and Ag content of electrum. Solid circle: epithermal Au-Ag vein-type deposits. Open circle: epithermaI base metal vein-type deposits. Solid square: hypo/mesothermal polymetallic vein-type deposits. Open square: epithermal Au disseminated-type deposits. I: Tada, 2: Toyoha, 3: Omidani, 4: Innai, 5: Ikuno, ~" qe-Inakuraishi, 7: Nebazawa, 8: Kawazu, 9: Todoroki, 10: Yatani, 11: Seigoshi, 12: Sado, 13: Takeno, 14.145 ,,awaji, 15: Yugashima, 16: Takadama, 17: Handa, 18: Konomai, 19: Sakoshi-Odomari, 20: Toi, 21: Sanru, 22: Arakawa, 23: Taio, 24: Chitose, 25: Hokuryu, 26: Okuchi, 27: Fuke, 28: Yamagano, 29: Akeshi, 30: Kasuga (Sbikazono, 1986).
in Japan is generally higher than that for epithermal Au-Ag vein-type deposits in Japan (section 1.4.3). It is also inferred that the deposition of argentite occurs by a pH increase, considering the following reaction. 2AgC12 + H2S --+ Ag2S + 4C1- + 2 H +
(1-48)
Therefore, it is likely that Ag-rich electrum and large amounts of sulfide minerals including argentite could precipitate due to CO2 loss and pH increase under low f s 2 and fo2 conditions. Therefore, this mechanism (boiling and gas loss from the hydrothermal solution with different fs2, J%2, CO2 concentration and pH) could explain why the Ag content of electrum correlates with Ag/Au total production ratio (Fig. 1.124). The above mechanism (boiling, loss of CO2 and increase in pH) could also lead to the deposition of other sulfides. The reactions causing sulfide depositions by this mechanism are written as, ZnC12 + H2S -+ ZnS + 2 CI + 2 H +
(1-49)
PbC12 +H2S --+ PbS + 2C1- + 2 H +
(1-50)
CuC1 + 2H2S + FeC12 --+ CuFeS2 + 3 C1- 4- 1/2 H2 + 3 H +
(1-51)
Miocene-Pliocene Hydrothermal Ore Deposits
173
As noted already, Ag content of electrum from the epithermal Au-Ag vein-type deposits is higher than that of the electrum from Kuroko deposits, and Ag/Au total production ratio of epithermal Au-Ag vein-type deposits (average 18) is lower than that of Kuroko deposits (average 76). Therefore, this relation is different from that found in the epithermal vein-type deposits. Ag/Au ratio of electrum may be controlled by the following reaction (Shikazono, 1981): (AU)el + 2H2S + Age12 = (Ag)el + 2 e l - + Au(HS)2 + 2 H +
(1-52)
It is thought from this reaction, that Au-rich electrum precipitates from ore fluids with high C1- concentration and low pH. Therefore, it is considered that different C1concentration and pH are important factors causing different relationship between Ag content of electrum and Ag/Au total production ratio of Kuroko deposits and epithermal vein-type deposits. Temperature is also important factor controlling Ag content of electrum. Shikazono (1981) showed that Ag content of electrum increases with decreasing temperature, assuming that EAg/EAu in ore fluids is constant, and dominant Au and Ag species are Au (HS)2 and AgC12. Figure 1.102 shows the iso-Au concentration contours (solubility of Au) on log fo2-pH diagram (Henley, 1984). The f Q - p H region for maximum Au concentration is close to H2S/HS- boundary and SO2-/reduced sulfur species (HzS, HS-) boundary. The most probable /O2-pH region of epithermal Au-Ag vein formation is also close to SO2-/reduced sulfur species boundary and the pH in equilibrium with adularia/sericite boundary, but the pH lower than adularia/sericite boundary is estimated based on the minerals coexisting with electrum: (1) adularia is abundant in the epithermal Au-Ag vein-type deposits but not in direct contact with electrum; (2) the minerals in direct contact with electrum are quartz, sericite/smectite and chlorite/smectite interstratified clay minerals. The most likely explanation for the variation of f Q - p H region of different epithermal Au deposition is the mixing of ascending hydrothermal solution whose foz-pH region is close to the region showing maximum Au solubility and descending acid sulfate hot spring whose temperature is lower than that of ascending hydrothermal solution. Acid sulfate solution contains appreciable amounts of SO]- but no H2S. This mechanism can explain the formation of Te-bearing Au-Ag veins in which sulfides are poor in amounts. The deposition of sulfides is generally difficult by this mechanism because solubilities of sulfides generally increase with decreasing of pH. However, if temperature of mixed fluid decreases considerably by this mechanism, the deposition of sulfides may be possible, because solubilities of sulfides due to chloro complexes decrease with decreasing of temperature. The above interpretation of formation of epithermal Au-Ag vein-type deposits is supported by (1) thermochemical calculations on this type of mixing (Reed and Spycher, 1985), and (2) the geological occurrence of epithermal Au-Ag vein-type deposits and associated advanced argillic alteration. Reed and Spycher (1985) pointed out based on thermochemical calculations that gold does not precipitate due to boiling because pH increases by boiling, leading to an increase in gold solubility due to gold-thio complexes. They indicated that acidification
Chapter 1
174
,,"0 ~
.,
Y . ~,~ --: [ , 2
I
f~)
J
~
," ',.,, Advanced Argillic ," /~ ,' "~" Alteration , ....,_Funabara area ,,s ~
"Seigoshi
".
-- ,~'~oi ,~l~ ~ ) m i n e , mine ~ ~. ,,1 '/
,"
,"
,"
', J"
"",
,',
F~ "
, ,,.
t
-
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-
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Proov -"
tic
:. mine
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&
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Mochikoshi
,, Ugusu mine .
~_.J
"
mine -
'
,'
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Diorite Porphyry Silicified Rock of Advanced Argillic Alteration r - ~ Gold-SilverQuartz Vein Fault
Figure 1.125. Distribution of epithermal Au-Ag vein-type deposits, propyIitic and advanced argillic alterations and intrusive rocks of diorite porphyry (Shikazono, 1985a).
of gold-bearing boiled waters by descending acid-sulfate waters is possibly of great significance to near-surface gold precipitation. The acidification destroys the dominant gold complex, Au(HS) 2, forcing gold precipitation by: H + + Au(HS)2 + 1/2 H2 ~ Au + 2 H2S
(1-53)
Shikazono (1985a) has studied hydrothermal alterations in the epithermal AuAg mine district in Izu Peninsula, middle part of Honshu, and indicated that (1) the propylitic alteration occurs widely in the district; (2) at the centre of the district and stratigraphicalty upper horizon, there exists advanced argillic alteration; (3) epithermal Au-Ag vein-type deposits are distributed at marginal zone in the district (Fig. 1.125); (4) the age of formations of advanced argillic alteration and epithermal Au-Ag veins are nearly the same (Au-Ag vein: 1.8-1.1 Ma; advanced argillic alteration: 2.2-1.2 Ma); and (5) epithermal Au-Ag deposition (Seigoshi) occurred in intermediale f Q - p H region of advanced argillic and propylitic alterations.
Miocene-Pliocene Hydrothermal Ore Deposits
175
By considering the features of hydrothermal system associated with epithermal Au-Ag mineralization, a model for the hydrothermal system is constructed in Fig. 1.123. Schematic diagrams showing the sites of Te-bearing and Se-bearing Au-Ag vein-type deposits in hydrothermal system are constructed based on the mineralogic, geologic and geochemical features of these deposits (Fig. 1.123). SO ] . in the acid sulfate hot springs responsible for advanced argillic alteration and/or epithermal Au-Ag ore fluids is considered to be of volcanic SO2 gas origin because ~348 of sulfates (alunite, barite) in advanced argillic alteration zone and in Te-bearing veins is high, more than +20%0. This is reasonably explained by the generation of SO ] - and H2S by the following hydrolysis reaction of volcanic SO2 gas. 4802 + 4 H20 "-'+ 3 H2SO4 + H2S
(1-54)
It seems likely that the mixing of acid sulfate solution with nearly neutral low salinity Au-bearing fluids seems the most likely mechanism for the formation of epithermal Au-Ag vein-type deposits. This mechanism as a main cause for epithermal-type Au deposition is supported by sulfur isotopic data on sulfides. Shikazono and Shimazaki (1985) determined sulfur isotopic compositions of sulfide minerals from the Zn-Pb and Au-Ag veins of the Yatani deposits which occur in the Green tuff region. The values for Zn-Pb veins and Au-Ag veins are ca. +0.5%0 to +4.5%o and ca. +3%o to +6%0, respectively (Fig. 1.126). This difference in 334S of Zn-Pb veins and Au-Ag veins is difficult to explain by the equilibrium isotopic fractionation between aqueous reduced sulfur species and oxidized sulfur species at the site of ore deposition. The non-equilibrium rapid mixing of H2S-rich fluid (deep fluid) with SOl--rich acid fluid (shallow fluid) is the most likely process for the cause of this difference (Fig. 1.127). This fluids mixing can also explain the higher oxidation state of Au-Ag ore fluid and lower oxidation state of Zn-Pb ore fluid. Deposition of gold occurs by this mechanism but not by oxidation of HzS-rich fluid. This type of mixing could reasonably explain the occurrence of acidic alteration minerals such as kaolinite and alunite in the low-sulfidation epithermal gold vein district (e.g., Seta in northeast Hokkaido, Hishikari in southern Kyushu) (Yajima et al., 1997) Mixing of high temperature hydrothermal solution with high salinity and low temperature solution with low salinity of meteoric water origin seems the most likely mechanism for the base-metal vein-type deposition. The salinity-temperature relationship obtained by fluid inclusion study on the Toyoha and Ohe P b - Z n - M n vein-type deposits and Fujigatani-Kiwada W skarn-type deposits indicates positive correlation (Fig. 1.128). This indicates that the mixing of high temperature fluids of deep-seated origin with low temperature fluid of meteoric water origin, together with the conductive cooling of mixed fluids caused the formation of these ore deposits (Fig. 1.129) (Shibue, 1991). Another possible mechanism for the ore deposition is boiling of fluids. For example, the solubility of galena is controlled by, PbS + 2H + + 2C1- = PbC12 + H2S
(1-55)
If boiling took place, H2S gas removes from the fluids. But CO2 loss causes an increase in pH, leading to the deposition of galena. However, no evidence of boiling of
Chapter 1
176
Galena
[]
Au-Ag
•
Zn-Pb vein
I
I
I
ll,ll
vein
i
,P1
,I"1 i
Chalcopyri le
Sphalerite
!
I
IL
~ 1
I
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i
~
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in
i
i
|
l
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I .,.I,,. ,
n
2
3
,
0
I
,nR
•B
, J
4
5
6 (~ 34 S ( ° / o o
)
Figure 1.126. Sulfur isotopic compositions of sulfide minerals from Au-Ag veins and Zn-Pb veins from the Yatani mine (Shikazono and Shimazaki, 1985).
fluids in inclusion fluids from the base-metal vein-type deposits is observed. Therefore, boiling seems unlikely as a depositional mechanism for galena as well as other common sulfides (chalcopyrite, sphalerite, pyrite).
1.4.5.2. Origin of ore fluids 3D and ~180 data on fluid inclusions and minerals at main stage of epithermal A u A g mineralization clearly indicate that the dominant source of ore fluids is meteoric water. Meteoric water penetrates downwards and is heated by the country rocks a n d / o r intrusive rocks. The heated water interacts with country rocks a n d / o r intrusive rocks and extracts sulfur, Au, A g and other soft cations (e.g., Hg, T1) from these rocks. If hydrothermal solution boils, it becomes neutral or slightly alkaline, leading to the selective leaching of soft cations such as Au, Ag, Hg and T1 from country rocks. However, a contribution of sulfur gas and other components from m a g m a cannot be ruled out.
Miocene-Pliocene Hydrothermal Ore Deposits
177
Au.Ag. ore fluid "~Zn.Pb
(a)
ore fluid
ore fluid
A
~-Zn.Pb
(b) 8 a4Sr.s.s.
ore fluid
D 8 34Sr.s.s.
Figure 1.127. (a) Schematic representation of change in ~SO42-/EH2S ratio (concentrationratio of oxidized sulfur species and reduced sulfur species) and 334S value of reduced sulfur species (334Sr.s.s.) in ore fluids during oxidation of reduced sulfur species at constant temperature and ES (total dissolved sulfur concentration). Dotted area represents the possible region for the fluids accompanied by oxidation of reduced sulfur species. A: initial values of 334Sr.s.s. and zsoZ-/EH2S in fluids prior to the oxidation. B: final values of 834Sr.s.s. and ~SO2 /EH2S in fluids when ~)34Sr.s.s. remains constant during oxidation. C: final values of 834Sr.s.s. and ESO2 /EHzS in fluids when isotopic equilibriumbetween reduced sulfur species and oxidized sulfur species is attained during the oxidation. Hatched areas show 334Sr.s.s. and ZSO]-/EH2S estimated for Zn-Pb ore fluid and Au-Ag ore fluid. (b) Schematic representation of change in ESO]-/EH2S ratio of ore fluids and 334Sr.s.s. value in ore fluids during mixing of t-I2S-richfluid (D) and SO42--rich fluid (E) at constant temperature. Dotted area represents the probable region resulted from the mixing of two fluids (D and E). F: final values of 834Sr~.s. and ESO]-/EH2S of the mixture of C and D fluids when isotopic equilibriumis attained betweenreduced sulfur species and oxidized sulfur species during the mixing. Hatched areas represent ~34Sr.s.s. and ESO]-/EH2S estimated for Zn-Pb ore fluid and Au-Ag ore fluid (Shikazono and Shimazaki, 1985).
8180 of late-stage hydrothermal solution is high (0%o to +3%o), as recognized in the Seigoshi and Hishikari A u - A g veins (Shikazono, 1988a; Shikazono and Nagayama, 1993). This increase in 8180 with the stage of hydrothermal system may be due to the change in water/rock ratio, boiling and kind of rocks interacting with fluids. ~D and 3180 data on fluid inclusions and minerals, 313C of carbonates, salinity of inclusion fluids together with the kind of host rocks indicate that the interaction of meteoric water and evolved seawater with volcanic and sedimentary rocks are important causes for the formation of ore fluids responsible for the base-metal vein-type deposits. High salinity-hydrothermal solution tends to leach hard cations (base metals, Fe, Mn) from the country rocks. Boiling may be also the cause of high salinity of base-metal ore fluids. However, this alone cannot cause very high salinity. Probably the other processes such as ion filtration by clay minerals and dissolution of halite have to be considered, but no detailed studies on these processes have been carried out. Origin of sulfide sulfur of epithermal base-metal veins is thought to be same as that of Kuroko deposits because average ~34S value of base-metal vein-type deposits is +4.7%o which is identical to that of Kuroko deposits (+4.6%o) (Shikazono, 1987b). Namely, sulfide sulfur of base-metal veins came from igneous rocks, sulfate of trapped seawater in marine sedimentary rocks, calcium sulfate (anhydrite, gypsum) and pyrite. ~34S of sulfide sulfur of epithermal base-metal vein-type deposits can be explained by the interaction of seawater (or evolved seawater) with volcanic rocks. There are the following three possibilities for the origin of sulfide sulfur of epithermal A u - A g vein-type deposits:
Chapter 1
178 t-" Q)
5-
a3
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Q)
o
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tco
0 100
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l
I
150
200
250
3O0
350
Homogenization (°C) t-8-
Ohe ~r
o
6-
sal-o e~.at. (198oI
o
o early
Z
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o~ c-
lafe •
2o 100
co ~Q)
o
04
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./.
ii 4ko4 •
•
•
[
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150
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250
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Homogenization (°C) 10-
Fujigaf'ani-Kiwada
>
5
8
0
6
•
•
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z,~
:.
i ||~'!% '°z ;• " .:. r'.'i ..
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°.~.:°.s• • **• %%* •• " % % . • o,~, • • • *o*e *'l* °
.
e-
co
0 loo
150
200
250
300
Homogenization (°C) Figure 1.128. Plots of homogenization temperature against salinity of fluid inclusions from the •he (Mn-PbZn), Toy•ha (Pb-Zn-Mn), and Fujigatani-Kiwada (W) deposits (Shibue, 1991).
(1) Leaching of sulfide sulfur from subaerial young (Miocene-Pliocene) volcanic rocks (Shikazono, 1987b). (2) Same origin of sulfur for base-metal vein-type deposits and Kuroko deposits. (3) Contributions both from sulfide sulfur leached from volcanic rocks and marine sulfate in Green tuff.
Miocene-t'liocene Hydrothermal Ore Deposits
179
c-
"-~
o - - o C o r r e t a f i o n I.ine
6
OQ
•
Conductive tooting 0
~0
0 0 ~
z 0
•~ co
0
1
100
150
0
000
I
O
I
200 250 Homogenization temperature (°C)
I
I
300
350
Figure 1.129. Illustrationof the regressionprocedureof fluid inclusion data on the Ohe deposit(Shibue, 1991).
Possibility (1) was proposed by Shikazono (1987b) who considered that the lower 3348 values of sulfide sulfur than base-metal vein-type deposits and Kuroko deposits can be explained by the leaching of sulfide sulfur from volcanic rocks with lower 334S values (0%0 to +5%o) (Uyeda and Sakai, 1984). (2) is also possible. It has been shown that a large amount of SO ] - together with reduced sulfur species such as H2S and H S - were present in epithermal Au-Ag ore fluids. Therefore, it is likely that 334S of reduced sulfur species whose values are typical Green tuff value decreases due to SO]--H2S(HS - ) fractionation. As noted already, Shikazono (1999b) divided epithermal Au-Ag deposits into Green tuff-type and Non-Green tuff-type based on their distributions. 334S values of Green tuff-type are higher than those of Non-Green tuff-type (Fig. 1.110). Host rocks, basement rocks, distribution of ore deposits, total production of Au and Ag, Ag/Au total production ratio, metals produced, 313C and 3180 of carbonates from the two types of deposits are different (Table 1.18). For instance, 334S values of sulfides from Green tuff-type are higher than those from Non-Green tuff-type. These differences are clearly interpreted by that sulfide sulfur of Green tuff-type was contributed by marine sulfate sulfur but that of Non-Green tuff-type is by igneous and sedimentary sulfide sulfur. The above-mentioned consideration on the origin of Neogene epithermal vein-type and Kuroko deposits is roughly consistent with the view by Mosier et al. (1986). They examined grades, tonnages and basement rocks for 88 epithermal precious and base-metal quartz-adularia-type districts in North and Central America and Japan and revealed that the type of basement rock below the mineralized veins is useful for predicting grade and size of deposits. Epithermal districts overlying basement with salt and evaporites of rocks with trapped sea water have a median tonnage (production and reserves) of 1.4 million metric tons and median grades of 1.5 g/ton Au, 130 g/ton Ag, 2.5% Pb, 1.7% Zn and 0.16% Cu, and districts overlying sedimentay basements have a median tonnage of 0.77 million metric tons and median grades of 7.5 g/ton Au, 110 g/ton Ag, less than 0.025% Zn, less than 0.005% Cu and less than 0.001% Pb, and districts overlying igneous basements have a median tonnage of 0.3 million metric tons and median grades of 5.9 g/ton Au, 38 g/ton Ag, tess than 0.25% Zn, less than 0.002% Cu, and less than
180
Chapter 1
0.003% Pb. Their results clearly demonstrate that basement rocks affect vein components in epithermal precious and base-metal quartz-adularia-type deposits. As noted already, epithermal vein-type deposits are classified primarily on the basis of their major ore-metals (Cu, Pb, Zn, Mn, Au and Ag) into the gold-silver-type and the base-metal-type. Major and accessory ore-metals from major vein-type deposits in Japan were examined in order to assess the possible differences in the metal ratios in these two types of deposits (Shikazono and Shimizu, 1992). Characteristic major ore-metals are Au, Ag, Te, Se and Cu for the A u - A g deposits, and Pb, Zn, Mn, Cu and Ag for the base-metal deposits (Shikazono, 1986). Accessary metals are Cd, Hg, T1, Sb and As for the A u - A g deposits and In, Ga, Bi, As, Sb, W and Sn for the base-metal deposits (Table 1.22, Shikazono and Shimizu, 1992). Minerals containing Cu, Ag, Sb and As are common in both types of deposits. They are thus not included in Table 1.22. As already noted in section 1.4.3, geochemical features of ore fluids responsible for base-metal and gold-silver types of deposits are distinct. They are summarized in Table 1.22. The differences in metals concentrated to the deposits and geochemical fectures of ore fluids responsible for both types of deposits are interpreted in terms of HSAB (hard, soft, acids and bases) principle by Pearson (1963, 1968) below. Ahrland et al. (1958) classified a number of Lewis acids as of (a) or (b) type based on the relative affinities for various ions of the ligand atoms. The sequence of stability of complexes is different for classes (a) and (b). With acceptor metal ions of class (a), the affinities of the halide ions lie in the sequence F - > CI- > Br- > I - , whereas with class (b), the sequence is F - < C1- < B r - < I - . Pearson (1963, 1968) classified acids and bases as hard (class (a)), soft (class (b)) and borderline (Table 1.23). Class (a) acids prefer to link with hard bases, whereas class (b) acids prefer soft bases. Yamada and Tanaka (1975) proposed a softness parameter of metal ions, on the basis of the parameters En (electron donor constant) and H (basicity constant) given by Edwards (1954) (Table 1.24). The softness parameter ~ is given by c~/(oe +/~), where c~ and/~ are constants characteristic of metal ions. They indicated that the softness parameter may reasonably be considered as a quantitative measure of the softness of metal ions and is consistent with the HSAB principle by Pearson (1963, 1968). Wood et al. (1987) have shown experimentally that the relative solubilities of the metals in H20-NaC1-CO2 solutions from 200°C to 350°C are consistent with the HSAB principle; in chloride-poor solutions, the soft ions Au + and Ag + prefer to combine with the soft bisulfide ligand; the borderline ions Fe 2+, Zn 2+, Pb 2+, Sb 3+ and Bi 3+ prefer water, hydroxyl, carbonate or bicarbonate ligands, and the extremely hard Mo 6+ bonds only to the hard anions O H - and 0 2 - . Tables 1.23 and 1.24 show the classification of metals and ligands according to the HSAB principle of Ahrland et al. (1958), Pearson (1963, 1968) (Table 1.23) and softness parameter of Yamada and Tanaka (1975) (Table 1.24). Comparison of Table 1.22 with Tables 1.23 and 1.24 makes it evident that the metals associated with the gold-silver deposits have a relatively soft character, whereas those associated with the base-metal deposits have a relatively hard (or borderline) character. For example, metals that tend to form hard acids (Mn 2+, Ga 3+, In 3+, Fe 3+, Sn 4+, MoO 3+, WO 4+, CO2) and borderline acids (Fe 2+, Zn 2+, Pb 2+, Sb 3+) are enriched in the base-metal deposits, whereas metals that tend to form soft acids
Miocene-Pliocene Hydrothermal Ore Deposits
181
TABLE 1.22 Accessory metals from vein-type deposits in Japan (Shikazono and Shimizu, 1992) Mine
Metal Hg
T1
Cd
Bi
Gold-silver type Konomai Kitanoou Showa Meiji Takadama Hirukodate Okuchi Hazami Osorezan Tsugu Chitose Todoroki Seigoshi Kushikino Teine Nishizawa Kawazu Ifiki Yatani Base-metal type Jizo Akenobe Toyoha Ashio Goka Miyatamata Agenosawa Akarimata Hosokura Nakanosawa Fukoku Ikuno Taishu Suttu Hayakawa Akagane Kishu Kutosan Kurokawa Goka Omidani Imaiishizaki Inakuraishi Tada Fukoku Omodani
Konjo Ryujima
Mo
Sn
W
X X
X
X X
X
X
X
X
X
X
X
X
X
X X X X X X X X X
X
Chapterl
182 TABLE 1.23 Classification of metals and Iigands according to the HSAB principle Mn2+, Ga3+, In3+, Co2+, Fe3+, As3+, Sn4+' MoO3+, WO4+, Co2+ Cu+ , Ag+ , Au+ , T1+, Hg+, Cd2+, Hg2+, Te4+ , TI3+ Fe2+, Co2+, Ni2+, Cu2+, Zn2+, pb2+, Sn4+, Sb3+, Bi3+, SO2 OH , CI-, CO2, SO2 H2S, HS-, S 2
Hard acids: Soft acids: Borderline acids: Hard bases: Soft bases:
TABLE 1.24 Softness parameter of the various metal ions Metal ions
Softness parameter
Metal ions
Softness parameter
Ag+ Hg+ TI+ Cu+ Cd2~ Ni2+ Zn3+ Co2+ Zn2+
1.03 1.01 0.98 0.96 0.96 0.94 0.93 0.92 0.91
CU 2+
0.89 0.87 0.85 0.84 0.82 0.78 0.73 0.58
Bi3+ Pb2+ Fez+ Mn2+ Fe3+ Sn2+ Ga3+
(Ag +, Au +, T1+, Te 4+, T13+) are enriched in the g o l d - s i l v e r deposits. Metals that have high values of the softness parameter (Ag +, Hg +, T1+, Cd 2+) are associated with the g o l d - s i l v e r deposits, whereas those that have low values o f the softness parameter (Zn 2+, In 3+, Bi 3+, Pb 2+, Te 4+, Mn 2+, Sn 4+, Ga 3+) are found with the base-metal deposits. These correlations mean that the HSAB principle could be a useful approach to evaluate the geochemical behavior o f metals and ligands in ore fluids responsible for the formation of the epithermal vein-type deposits. A m o n g the ligands in the ore fluids, H S - and H2S are the most likely to form complexes with the metals concentrated in the gold-silver deposits (e.g., Au, Ag, Cu, Hg, TI, Cd), whereas C1- prefers to form complexes with the metals concentrated in the base-metal deposits (e.g., Pb, Zn, Mn, Fe, Cu, and Sn) (Crerar et al., 1985). Generally, the complexes with intermediate or hard ligands (e.g., chloro complexes) should become more stable with increasing temperature than complexes with soft ligands (e.g., bisulfide complexes) (Seward, 1981 ; Crerar et al., 1985). The higher temperatures of formation of the base-metal deposits (Table 1.13) also are in accordance with the HSAB principle. It was shown in Table 1.13 that the base metal and g o l d - s i l v e r types of deposits formed at different temperatures and concentrations of C I - , H S - , and CO2. Thus, it could further be inferred that the H S A B principle can be successfully applied to the genesis of these vein-type ore deposits, formed mostly at less than ca. 250°C. Crerar et al. (1985) noted that Pearson's rule (the HSAB principle) successfully describes speciation to about 250°C, but may break down at higher temperatures, as all metals become harder.
Miocene-Pliocene Hydrothermal Ore Deposits
183
A few appilications of the HSAB principle to hydrothermal ore deposits have been carried out (Crerar et al., 1985; Wood et al., 1987; Shikazono and Shimizu, 1992). These studies demonstrate that the HSAB principle is useful in interpretations of the metal ratios in ore deposits. For example, Wood et al. (1987) has shown that gold is transported by lower salinity fluids than the base metals, and this difference in salinity is a significant factor in the separation of gold and base metals in Archean deposits in greenstone belts. Cathles (1986) also has indicated that the solubility of gold is much greater in lowsalinity solutions; this can explain the bimodal populations of base-metal-rich, gold-poor deposits (stratiform deposits similar to Kuroko deposits in Japan) and base-metal-poor, and gold-rich deposits in greenstone belts (lode gold deposits). Such differences in the salinity of ore fluids responsible for the epithermal gold deposits and Kuroko and base-metal vein-type deposits in Japan have been pointed out also by Shikazono and Shimizu (1993). Therefore, it could be inferred that the difference in salinity is a main cause for a separation of gold and base-metals in mineralized zones in young (Tertiary to Quaternary) Japanese epithermal vein-type and Kuroko deposits in volcanic terranes and in Archean deposits in greenstone terranes. However, the concentration of CO2 in ore fluids responsible for Japanese gold-silver deposits and Archean lode gold deposits in greenstone belt seems to be different; CO2 concentration is low (0.01-0.1 molal) for Japanese deposits (Shikazono, 1985a), whereas it is high (0.05-2 molal) for Archean deposits (Wood et al., 1987). The reason for this difference is not known.
1.4.6. Hishikari deposit: an example of Japanese epithermal Au-Ag vein-type deposits One of the most important steps made in the research of Japanese epithermal gold deposits during the last two decades is the comprehensive study of the Hishikari gold deposits which occur in southern Kyushu (Fig. 1.130). The deposits are characterized by high gold grade and large amounts of gold reserve. Average gold grade in the Honko (Main) deposits is ca. 70 g/metric ton which is enormously high, compared with the other Japanese epithermal gold deposits. Ore reserve is estimated to be about 250 metric tons Au, which is the largest among the Japanese epithermal gold deposits. Ag/Au production ratio is about 0.7, which is relatively low. The Hishikari deposit is only one which belongs to a giant (Laznicka, 1983) and bonanza (Sillitoe, 1993) epithermal-type deposit (Izawa, 2001). Much research on the Hishikari deposits has been carried out since the discovery of gold veins in 1981. Urashima and Izawa (1983) reported fluid inclusion studies using core samples. Abe et al. (1986) made a detailed description of the veins. The regional geology of this district is described in MMAJ and SMM (1987) and Urashima et al. (1987). Izawa et al. (1990) reviewed the studies on the geology, geophysics and geochemistry which had been done during 1980s. The Special Issue of Resource Geology on the Hishikari deposits (Shikazono et al., 1993) includes various aspects of the Hishikari deposits (oxygen isotopes of gangue minerals, hydrothermal alteration, precipitation sequence, fluid inclusions, vertical electric profiling and electric sounding surveys, structural geological analysis, opaque minerals,
Chapter 1
184
(/-f / -~ k~
~
k ) ,
[ | ]
• Hot springtype O EpithermalAu-Agvein-type A Activevolcano
.~v-/
o
. , ~ /
r
20 ,
I
40 ,
I(km)
00kuehi
/ Miyazaki0 (
OYamagano
Kagoshimao) t. A /.j
e /
131E Figure ].130. A map showing the location of the Hishikati mine.
prospecting, etc.) which had been studied during late 1980s and early 1990s. No review on the studies during the last decade has been published in English. Thus, a review of the studies done during the last decade and recent studies by the author will be summarized below.
1.4.6.1. Geology and vein system A detailed general geology of the Hishikari district has been done by many investigators (Abe et al., 1986; MMAJ and SMM, 1987; NEDO, 1991; Izawa et al., 1993). The general geology of this district is briefly described below. The district is composed of sedimentary rocks of the pre-Paleogene Shimanto Supergroup (dominantly shale and sandstone) and Quaternary andesitic and dacitic volcanic rocks. The Shimanto Supergroup is comprised of shale, sandstone and their alternations. Although no fossil data are available, the age of sedimentation is thought to be middle to upper Cretaceous age from its lithology (Izawa et al., 1990). The Shimanto
185
Miocene-Pliocene Hydrothermal Ore Deposits
Supergroup rocks in the Hishikari district suffered hydrothermal alteration. Chlorite, quartz and sericite occur abundantly near the veins. The other constituents are pyrite, albite, calcite and organic matter. Quaternary volcanic rocks unconformably overlie the Shimanto Supergroup. Quaternary volcanic rocks consist of Hishikari Lower Andesites (0.98-1.62 Ma), Hishikari Middle Andesites (0.78~0.79 Ma), Kurozonsan Dacites (0.95-1.56 Ma), Shishimano Dacites (0.66-1.6 Ma), and Hannyaji Welded Tuff (0.56-0.731 Ma) (Izawa et al., 1990, 1993). The Hishikari Lower Andesites consist of hypersthene-augite andesite lava flows in the upper horizon and andesitic pyroclastic rocks in the lower horizon. The Kurozonsan Dacites overlying the Hishikari Lower Andesites dominantly consist of hyperstheneaugite-bearing biotite-hornblende dacide lava flows. The Hishikari Middle Andesites overlying the Hishikari Lower Andesites consist of hypersthene-augite andesite lava flows and pyroclastic rocks. The Shishimano Dacites overlying the Hishikari Middle Andesites consist mainly of biotite-hornblende dacite lavas. The Hishikari Upper Andesites overlying the Hishikari Lower Andesites and the Shishimano Dacites consist of hypersthene-augite andesite lava flows and their pyroclastic rocks. The deposit consists of three vein systems, the Honko, Sanjin and Yamada veins (Fig. 1.131). The veins strike N30°E to N50°E and dip 70-90°NW. The veins are hosted
0
0.5
I
,
1 km k, I
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25
. . . .
, ,. __,, " - ~ ' ~
J
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25
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I1:zoneII II1: zone III
Figure 1.131. Plan of alteration mineral zonation of the volcanic rocks in the Hishikari area (Izawa et al., 1990).
Chapter 1
186
in hydrothermally altered Shimanto Supergroup (Honko (Main) and Sanjin veins), and Quaternary andesite (Yamada veins).
1.4.6.2. Hydrothermal alteration Izawa et al. (1990) recognized the following alteration zones from the vein towards margin of the Hishikari A u - A g mine area, chlorite-sericite zone (zone IV), interstratified clay mineral zone (zone III), quartz-smectite zone (zone II) and cristobalite-smectite zone (zone I) and least altered zone (L.A. (least altered) zone) (Fig. 1.131).
1.4.6.3. Mineralogy Quartz and adularia are dominant gangue minerals in the veins. Small amounts of sericite/smectite, calcite, and truscottite occur as vein staff. Fine-grained electrum, naumannite, Ag-sulfosalts (pyrargyrite, polybasite), chalcopyrite, galena, pyrite, marcasite and stibnite are found. Detailed mineralogical descriptions are given in Urashima and Izawa (1983), Urashima et al. (1987), Izawa et al. (1990), and Nagayama (1993a). The sequence of mineralization has been studied by Nagayama (1993a) and Takahashi et al. (1998) (Fig. 1.132). Quartz is the most abundant mineral occurring throughout the vein. Adularia tends to occur at earlier stage. Smectite is the earliest mineral. Electrum tends to occur in early and middle stages. Naumannite occurs after the early-stage
"•,•quenceveil,of
I
t
a
miner~,.,
b
c
d
e
II f
g
h
111
i
Quartz
m
m
Adularia -i
~0
Smectite Electrum &"
.k"
:6_'
e6_
Naumannite
Ag-Sb-S "E ©
minerals
Pyrite Chalcopyrite Sphalerite m
abundant ~
rich
~
common-
poor - - f e w
Figure 1.132. Paragenetic sequence of the Hosen No. 5 vein (Takahashiet ai., 1998).
Miocene-Pliocene Hydrothermal Ore Deposits
187
electrum. Ag-Sb-S minerals are found both in early- and late-stages. Abundant pyrite occurs in early-stage. Small amounts of chalcopyrite are widely distributed. Sphalerite is found in late-stage. Stibnite and realgar are found in the latest stage. The sequence of mineralization from early to late is consistent with that of the other deposits (Asahi, Ogane, Takadama) sulfide stage (pyrite, marcasite, chalcopyrite, sphalerite, galena, electrum, naumannite, polybasitc, quartz, adularia), sulfosalt stage (pyrite, chalcopyrite, electrum, tetrahedrite, pyrargyrite, miargyrite) adularia, quartz (calcite), and barren stage (adularia, quartz, pyrite, realgar, stibnite) (Fig. 1.133). Adularia/quartz ratio decreases with the stage of mineralization and correlates with Au content of ore (Shikazono and Nagayama, 1993).
1.4.6.4. Geochemicalfeatures Oxygen and hydrogen isotopes. Oxygen and hydrogen isotopic studies on the altered rocks and gangue minerals were reported by Shikazono and Nagayama (1993), Naito et al. (1993), Matsuhisa and Aoki (1994), and Imai et al. (1998). Shikazono and Nagayama (1993) found that 8180 of adularia and quartz and adularia/quartz ratio decrease from early- to late-stage of the Hishikari mineralization, gold contents of ore samples in a vein positively correlate with 3180 and adularia/quartz ratio, as well as, the Pb, Cu, Ag and Se contents of vein samples, and Hg and As contents of the same samples increase towards the late-stage of mineralization (Fig. 1.104). Using fluid inclusion data and 3180 of adularia and quartz, it is inferred that 8180 of ore fluids decreases with the stage of mineralization (Shikazono and Nagayama, 1993). The early-stage ore fluids were of highly exchanged meteoric waters or magmatic water origin, while late-stage ore fluids were dominated by meteoric water (Shikazono and Nagayama, 1993). 8180 and 3D of ore fluids were influenced by clay minerals in the Shimanto Supergroup sedimentary rocks (Imai et al., 1998). The 813C value of H2CO3 and the 8~80 value of the hydrothermal solution using the data on vein calcite coexisting with electrum and assuming isotopic equilibrium with calcite and the temperature obtained by fluid inclusion microthermometry range from -14.4%o to -9.1%o, and from -6.2%o to +5.4%o, respectively, suggesting that the hydrothermal solutions isotopically equilibrated with the sedimentary basement rocks were responsible for the gold mineralization associated with calcite (Imai and Uto, 2001). Naito et al. (1993) showed that oxygen isotopic composition (81SO) of altered volcanic rocks in the Hishikari mine area varies systematically; +5.9%o to +15.9%o (zone I), +7.1%o to +12.4 %0 (zone II), +2.8 to +11.7 %o (zone III), and +2.1%o to +8.2 %o (zone IV) (Fig. 1.134). They calculated the change in 8280 values of hydrothermally altered volcanic rocks as a function of water to rock ratio by weight and temperature, assuming that oxygen isotopic equilibrium is attained in a closed system, and demonstrated that the increase in 8180 values of altered andesitic rocks from the veins towards peripheral zones can be interpreted as a decrease in temperature from the vein system (Fig. 1.135). In their calculations, the effect of mixing of hydrothermal solution with groundwater was not considered.
Chapter 1
! 88
Later
Earlier
Sulphide stage Sulphosalt stage Adularta Quartz Calcite Pyrite Marcasite Sphalerlte Chalcopyrlte Galena Gold Argentita Polybasite Tetrahedrlte Pyrargyrlte Mlargyrlte Realgar
mare
Barren stage
i
I
iwua=
I lib i
milm i
-? n
.,,?
Paragenetic sequence of the Asahi deposit Olta pref. Matsukuma (1951)
II Quartz Adularla Pyrite Chalcopyrlte Sphalerlte Galena Ag-mlnerele Gold Rhodochroalte Barite
~
III
-
~
--
-.-,
Psregenetic sequence of the Nishltanl vein, Nlahltanl vein group, the Ogane mine, Hokkaldo. Aklba (1957)
A
B
C
D
Quartz Adularla A:Big adularla crystals with fine quartz B:Aggregatee of adularla and quartz C:Masslve quartz with adulsrla l D:Drusy quartz General feature of paragenetlc sequence of gangue minerals, Tekatama mine, Fukushlma pref, modified from Yagyu (1954) Figure 1.133. "The normal order" in other epithermal gold deposits in Japan (Nagayama, 1993a).
Miocene-Pliocene Hydrothermal Ore Deposits
189
I 8.9+-1.4%-,n=50)
10
L.A. (least altered)
III IIII IIIIII
iI
Zone 110.9:1:2.7%.,n:~,l Cr-Sm 5 Cr-Ka(Ko)
6
llll I
i 1 i i I-I-I I11
I II
II
IK~KaI
5
lo
15
t I[ Zone (9.0±1.8%o, n=ll) Oz-Srn 5 Qz-Ko(Ko)
6
g
I"l R - F ~
Zone ( 6.9-+ 1,8% , n=32} 5 t Int.ClaYmirerolc
6
/ CM-Ser l |/
. o
I-13
I I III
5
I I I
Vl~ lo
VI
1~
I , r-I I-I
lo
1is
[7 Illllll [-'1 i i Ii Ii Ii Ii Ii 'i 'i I I I rl s
[-1
lb
1~
a~8o (%0) Figure I.i34. Histograms of 3[80 values of the volcanic rocks for five distinct alteration zones (Naito et aI., 1993). Shikazono et al. (2002) interpreted the ~180 zonation based on hydrothermal solution-groundwater mixing model. For the calculations the following assumptions were given: (1) Hydrothermal system at discharge zone is composed of five reservoirs such as ore deposit/zone IV boundary, zone III/II boundary, zone I I / I boundary, zone I/fresh country rocks boundary and temperature of each reservoir is 250°C, 220°C, 150°C, 100°C, and 25 °C, respectively.
Chapter 1
190
+20.C
250C-
+15. O
~
-
65°C
O 0 +10.0 ,.,f
100°(:;_ 113°C_
0
t43
1500C +5.0
200°C 250°(3 300°(3 I
i
i
I
i
1
I
1
2
3
4
5
6
7
9
Water/Rockratioin Weight Figure 1.I35. Change in 5180 values of hydrothermally altered rocks as a function of water to rock ratio with several different temperatures. The initial 3180 values of rock and water were taken as +8.5%o and -5.0%o, respectively (Naito et al., 1993).
(2) Initial 3180 of hydrothermal solution and groundwater are 0%0 and -7%0, respectively. (3) Oxygen isotopic equilibrium between mixed fluid and alteration minerals is attained. (4) Minerals in oxygen isotopic equilibrium with mixed fluid are feldspar for ore deposit/zone IV, and zone III/II, montmorillonite, and kaolinite for zone II/I, and montmorillonite for zone I/fresh rocks. Oxygen isotopic fractionation factors used for the calculation were taken from Taylor (1997). Initial 8180 value of hydrothermal solution (0%0) was estimated from 8180 values of K-feldspar and quartz in the veins and homogenization temperatures (Shikazono and Nagayama, 1993), and that of groundwater (-7%0) was estimated from meteoric water value of the south Kyushu district (-7%0) (Matsubaya et al., 1975). The mixing ratio of hydrothermal solution and groundwater was calculated based on the temperature of each reservoir. Using the estimated 81SO of fluid and oxygen isotopic fractionation between water and mineral, 3180 of altered rocks were estimated. The calculated result, together with the average 8180 of hydrothermal alteration zone by Naito et al. (1993), is shown in Fig. 1.136. This shows a fairly good agreement
Miocene-Pliocene HydrothermaI Ore Deposits
191
16
X 14
- • - Analytical value
12
~
--X-- Calculated value
.~ ~o
.....
o
£
• x. , -~~ . ~
,;o
,;o
=;o
=;o
aoo
Temp.(*(;)
Figure 1.136. 3180 change due to the mixing of hydrothermal solution and groundwater (Shikazono et al., 2002). between the model calculation and analytical data on 8180 except for the data on the most peripheral zone, suggesting the mixing of hydrothermal solution and groundwater is important for causing of 8180 variation in hydrothermally altered andesite.
~34S of sulfides. 834S values of sulfides in the veins are -1%o to +2%0 (Shikazono, unpublished). This is very close to average 834S value of Japanese epithermal A u - A g vein-type deposits which is ca. +2%o. In contrast to the 834S of hydrothermat solution for the vein, that of pyrite in hydrothermally altered rocks (Shimanto Shale) varies very widely, ranging from -5%o to +15%o. Based on the microscopic observation, pyrite with low 834S values less than 0%o is usually framboidal in form, suggesting that low 834S was caused by bacterial reduction of seawater sulfate. There are two possible interpretations of high 834S values (+10%o to 4-15%o). One is the reduction of seawater sulfate in a relatively closed system. The other one is a contribution of volcanic SO2 gas. As noted already, volcanic SO2 gas interacts with H 2 0 to form H2SO4 and H2S. 834S value of SO 2- formed by this reaction is generally high although the 834S value depends on the sulfur isotopic fractionation between H2SO4 and H2S and temperature. It is worth noting that in both cases a considerable degree of sulfate reduction took place. This may be due to the small water/rock ratio in the altered rock system compared to open crack systems through which large amounts of hydrothermal solution migrate. Bulk compositions of altered andesite. Bulk compositions of altered andesite were obtained by XRF analysis (X-ray fluorescence analysis) (Shikazono et al., 2002). Figures 1.137 and 1.138 show the change in elemental contents (SiO2, K20, CaO, MgO) of altered andesitic rocks away from the quartz vein system. Figure 1.139 shows the relationships between (CaO + Na20) content and K 2 0 content. These data indicate the following features of compositional variation of altered rocks.
I'o
90
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.
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260
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600
~00
D i s t a n c e f r o m the vein (m)
Figure 1.137, The variations of major element contents in andesite (drilling core, 8-MAHAK-4 and underground samples) away from the vein system (Shikazaoo et al., 2002), Diamond: Hishikari Lower Andesitic tuff (underground samples): square: Hishikari Lower Andesitic tuff (drilling core samples); triangle: Hishikari Lower Andesitic lava (drilling core samples); ×: relatively fresh Hishikari Lower Andesitic lava (drilling core samples). (A) SiO2 content variation. (B) K20 content variation. (C) CaO content variation.
193
Miocene-Pliocene Hydrothermal Ore Deposits 4
3.5 3 o~ 2.5
*
|
2*
4'
•
0
if2 1.5 1 0.5 0
0
Distance from the vein (m)
Figure 1.138. MgO content variation for underground samples. Diamond: Yusen No. 7 vein; square: Seisen No. 8 vein.
14
12
10
1( "x ×
•
~8 o ll•
Z +
Abbreviations: A zone I, × zone II, , zone III, • zone tV, + z o n e IV ( n e a r the quartz vein), • fresh andesite (Hishikari Lower Andesite)
O6 o
AA
X &
2
• • • ÷
÷
•
+%
÷ #
÷
04.+~
++'~÷
÷
o
0
5 K20 (wt%)lO
15
Figure 1.139. The relationship between K20 content and (Na20 + C a • ) content in andesite (Shikazono et al., 2002).
Chapter I
194
(1) SiO2 content of andesite tends to be high near the veins. High SiO2 content may be due to the presence of quartz veinlets (Fig. 1.137A). (2) K20 and MgO contents of altered andesite decreases away from the vein (Fig. 1.137B and Fig. 1.138). (3) Analytical data on andesite are plotted on (CaO 4- N a 2 0 ) - K 2 0 diagram (Fig. 1.139). This shows that (CaO 4- Na20) content inversely correlates with K20 content. These results are consistent with XRD (X-ray diffraction) results. The amounts of K-feldspar, K-mica and chlorite are higher in the altered rocks closer to the veins and Ca-zeolites and smectite decrease in amounts towards periphery of the alteration zones. Although data are scattered and mostly no systematic correlation of Na20, CaO and MgO contents with 8t80 exist, K 2 0 content and 8~80 seem negatively correlated; K20 content is low and high for zone I and IV and 8180 is high and low for zone ! and IV, respectively. This correlation suggests that mineralogical alteration zoning correlates to 8180 variation that has been cited by Naito et al. (1993). The variations in K20, Na20 and CaO contents could be interpreted by thermodynamic consideration. The dependence of concentration of K +, Na +, Ca 2+ and H4SiO4 in equilibrium with common alteration minerals (K-feldspar, Na-feldspar, quartz) on temperature is shown in Fig. 1.140 (Shikazono, 1988b). This figure demonstrates that (1) chemical compositions of hydrothermal solution depend on alteration minerals, temperature and chloride concentration, and K + and H 4 S i O 4 concentrations increase and Ca 2+ concentration decrease with increasing of temperature. In this case, it is considered that potassic alteration adjacent to the gold-quartz veins occurs when hydrothermal solution initially in
E ..go
B
Na+
H
"
Ca2+
.A N
-1
-2
-3
1;0
200
2;0
300
Temp. (°C)
Figure 1.140. The dependence of concentration of K +, Na +, Ca 2+ and H4SiO4 in equilibrium with common alteration minerals (K-feldspar, Na-feldspar, quartz) with temperature (Shikazono, 1988b). Thermochemical data used for the calculations are from Helgeson (1969). Calculation method is given in Shikazono (1978a). Chloride concentration in hydrothermal solution is assumed to be 1 mol/kg H20. A-B: Na + concentration in solution in equilibrium with low albite and adularia. C-D: K + concentration in solution in equilibrium with low albite and adularia. E-F: H4SiO4 concentration in solution in equilibrium with quartz. G-H: Ca 2+ concentration in solution in equilibrium with low albite and anorthite.
Miocene-Pliocene Hydrothermal Ore Deposits
195
Total mKCJmHCI 600 0 500 o
400
Q. 300 E 200
Andalusite~ ~. ..... ~
K-feldspar
Pyrophylli~__ ~ ~
H~A
Kaolinite
100
Acidic GW 0
2 4 Total mKcJmHcI
GW I 6
Figure 1.141. Temperalure-aK+/aH+(a: activity) trend due to the mixing of hydrothermal solution and groundwater accompanied by hydrothermal alteration (Hemley and Jones, 1964). HS: hydrothermal solution, GW: groundwater, A: hydrothermaIsolution-groundwater mixing line, B: hydrothermal solution acidic groundwater mixing line (Shikazonoet al., 2002). equilibrium with propylitic alteration minerals (albite, K-feldspar, quartz) ascends rapidly and interacts with country rocks at lower temperature ( I - J - K in Fig. 1.140). It is also considered that K + is added from hydrothermal solution to the rocks as K-mineral such as K-feldspar and K-mica at the site of ore deposition, accompanied by the destruction of plagioclase in the country rocks and liberation of Ca 2+ and Na + to the fluid. It is expected that SiO2 content of the country rocks increases with proceeding of the alteration because solubility of SiO2 decreases with a decrease in temperature ( O - P - Q in Fig. 1.140). Figure 1.141 shows that the hydrothermal alteration zoning from K-feldspar through K-mica to kaolinite from the vein towards marginal part of alteration zones can be explained by a combined change of temperature and m~+/mH+ (m: molal concentration) ratio. Decreases in temperature and mK+/mH+ ratio of hydrothermal solution cause the above alteration zoning. This trend suggests that ascending hydrothermal solution mixed with acidic (low pH) and low-temperature descending groundwater (B in Fig. 1.141). Acidic groundwater may be formed by the interaction of groundwater with acidic alteration zone at shallow part, input of volcanic SO2 gas into groundwater, condensation of steam, or oxidation of H2S. If K + is added to the rock accompanied by the destruction of feldspar, the following reactions occur. CaO (Ca-feldspar) + 2 K + ~ K 2 0 (K-feldspar) + Ca 2+
(1-56)
Na20 (Na-feldspar) ÷ 2 K + --+ K 2 0 (K-feldspar) + 2 Na +
(1-57)
It is expected from (1-56) and (1-57) that CaO and Na20 contents of altered rocks inversely correlate with K20 content of altered rocks with a negative slope of - 1 . Analytical results show the negative correlation on (CaO + N a 2 0 ) - K 2 0 diagram
Chapter 1
196
(Fig. 1.139). However, the low CaO 4- Na20 content and K20 content data on Fig. 1.139 cannot be explained only by these reactions. The relatively low contents data could be attributed to the dissolution of silicates by acidic solution which was generated by the disproportionation reaction, 4SO2 (volcanic gas) + 4 H 2 0 ~ 6H + + 3 SO 2- + HzS
(1-58)
The following dissolution of silicates such as feldspar occur by the interaction of H + with feldspar, CaO (Ca-feldspar) + 2H + ~ Ca 2+ 4- H20
(1-59)
Na20 (Na-feldspar) 4- 2H + ~ 2Na + 4- H20
(1-60)
K20 (K-feldspar) 4- 2H +
(1-61)
--+ 2K + 4- H20
Therefore, it is thought that the mixing of acidic solution with hydrothermal solution occurred and andesite near the gold-quartz veins suffered superimposed potassic and advanced argillic alterations.
1.4.6.5. Interpretation of Si02 mineral zoning in terms of kinetics-fluidflow mixing model
Shikazono et al. (2002) considered the depositional mechanism of quartz and cristobalite and the change in silica concentration of fluid migrating through the altered rocks in the Hishikari mine district based on kinetics-fluid flow mixing model. Their discussion is summarized below. Hydrothermal solution containing appreciable amounts of dissolved silica migrates through andesitic volcanic rocks, accompanying SiO2 precipitation. Figure 1.142 shows the dependence of solubility of SiO2 minerals (quartz, cristobalite) on temperature. As described already, cristobalite occurs in peripheral and shallower part of hydrothermal alteration zone. Quartz is present in zones occurring in deeper and closer to the gold-quartz veins. Such zoning from quartz to cristobalite is also common in main active geothermal systems (Hayashi, 1973; Takeno et al., 2000). Precipitations of quartz and cristobalite occur due to a decrease in temperature that is caused by heat conduction, and fluid mixing. The changes in the concentration of dissolved silica during these processes are shown in Fig. 1.142. Decrease in temperature due to heat conduction alone cannot explain the distributions of quartz and cristobalite. The temperature at which the heat conduction trend crosses the cristobalite saturation curve is ca. 200°C, which is higher than the 100°C corresponding to the cristobalite/quartz boundary in active geothermal system (Hayashi, 1973; Takeno et al., 2000). The curve for the mixing of hydrothermal solution and groundwater always lies below the cristobalite saturation curve if the hydrothermal solution is in equilibrium with quartz. Therefore, the heat conduction and mixing of fluids in equilibrium with quartz are considered to be not main causes for the precipitation of quartz and cristobalite. Therefore, in order to know the change in dissolved silica concentration and temperature during the precipitation of quartz and cristobalite and mixing of fluids, the following equation could be used:
Miocene-Pliocene Hydrothermal Ore Deposits 800
197
800i
i (B)
(A) 700
700!
[]
E~ 0
o
600
6OO
[]
[3
~
0 H.S.--
500
400
[] H.ST"-
500
4O0
g o5
8 O300
300'
200 i
200
~
I 100
~.~.
/
._~ ~ 2
~
100 150 Temp (°C)
200
250
0.1
A / M ~ 1.00
l°° i ~ D D - ^ A
50
~/~
[3
o
%ee~%5
AA-
~
550
200
250
Temp (°C)
Figure 1.142. The computed result of the relationship between dissolved silica (H4SiO4) concentration of mixed fluid and temperature based on four reservoirs model (Shikazono et aI., 2002). Open triangle: solubility curve for quartz, Open square: solubility curve for c~-cristabalite,Solid triangle: Hishikari Lower Andesite lava (drilling core), Cross: Relatively fresh Hishikari Lower Andesite lava (drilling core). H.S.: hydrothermal solution; G.W.:ground water.
dC/dt = k(A/M)(Co - C) ÷ ( q l C 1 / V ) + (q2C2/V) - (ql + q2) C~ V
(1-62)
where C = concentration o f dissolved silica of output fluid ( m o l / k g H 2 0 ) , t = time (s), k = rate constant (mol m - 2 s 1), Co = concentration of dissolved silica saturated with respect to quartz ( m o l / k g H 2 0 ) , C1 = concentration of H4SiO4 o f input hydrothermal solution ( m o l / k g H 2 0 ) , C2 = concentration of H4SiO4 of input groundwater ( m o l / k g H 2 0 ) , ql = volume flow rate of hydrothermal solution (m 3/s), q2 = volume flow rate of groundwater (m3/s), M = mass of aqueous solution in a system (kg), V = volume of aqueous solution in a system (m3), and A = surface area of rocks which contacts with aqueous solution (m2). If steady state is attained, we obtain,
C = {k(a/M)Co + (ql C1 + q2C2/V } / { k ( a / M ) + (ql + q2)/V }
(1-63)
As shown in Fig. 1.143, the system is divided into four reservoirs. Each reservoir corresponding to alteration zone IV, III, II and I is assumed to be homogeneous with respect to temperature and concentrations of dissolved silica in aqueous solution. The temperature of the initial hydrothermal solution is assumed to be 250°C from homogenization temperature of fluid inclusions in vein quartz (Shikazono and Nagayama, 1993). Temperature o f each reservoir was estimated from the assemblage of hydrothermal alteration minerals and temperature of alteration zone in active geothermal system (e.g., Hayashi, 1973; Takeno et al., 2000).
Chapter 1
198 G°W.
H.S. ~ T ~250~
T ----220
T ~150
[ T ~-100
T =25
Figure 1.143. Four reservoirs model. H.S.: hydrothermal solution, G.W.: groundwater (Shikazono et at., 2002).
Volume flow rates of incoming hydrothermal solution and groundwater to each reservoir were estimated from the temperature of reservoirs and initial hydrothermal solution (250°C) and groundwater (25°C) Volume of solution (V) in a reservoir is expressed as, V = Vbox~b/100
(1-64)
where Vbox = volume of reservoir as a rock, and q~ = porosity. Volume flow rate (q) is expressed as, q = A'v(a/lO0
(1-65)
where A I = cross-section area of box (m 2) and v = velocity of fluids. It is assumed that the following simple relation is approximately established in a case of two fluids (fluid 1 and fluid 2) mixing. ql 7"i ÷qzT2 = (qI + q z ) T
(1-66)
where T = absolute temperature. The mixing ratio for fluid (R) is approximated as, R = q2/(ql + q2)
(1-67)
The calculations based on four reservoir models were made using equations (162)-(1-67) and precipitation rate constant (k) for SiO2 minerals by Rimstidt and Barnes (1980). Volume of each reservoir was calculated from the volume of each alteration zone and assuming porosity of alteration zone to be in a range of 3-1%. Dissolved silica concentration of groundwater was assumed to be 60 mg/kgH20 that is average dissolved silica concentration of groundwater in andesitic volcanic region in Japan (Shikazono, unpublished). Assuming the ranges of A / M , flow rate of mixed fluid, porosity and giving the precipitation rate of SiO2 minerals (Rimstidt and Barnes, 1980), the relationship between dissolved silica concentration of mixed fluid and temperature was obtained (Fig. 1.142). It was found that the porosity does not change the results of calculations. Figure 1.144 shows the results of calculation based on multireservoirs (40 reservoirs) model in which each reservoir corresponding to each alteration zone is divided into
Miocene-Pliocene Hydrothermal Ore Deposits
199
800
800
700
700 :
[] [2
600
E3 E3
600 []
[] C] H.S.-- ,
5O0
~ 400
[] H.S.~
400 [3 °
8
G
500
oo
G
b5 3OO
300 A/M~0.0
V = l O -3 ~
__10~.2
200
100200 G
~
I
°~'a
-4¢dM~1.00
1O0
50
100
150
Temp (°C)
200
250
50
100 150 Temp (°C)
200
250
Figure t. 144. Computedresults for multireservoirs (40) model. Open triangle: solubility curve for quartz, Open square: solubility curve for c~-cristabalite, Cross: no precipitation, Open circle: computedresult (Shikazono et al., 2002). ten reservoirs and temperature of each reservoir was given. The results of calculation indicate that the flow rate, 10.4.2 m/s, is the best estimate for A/M = 0.1 that is estimated from the width of quartz veinlets in altered andesite, and this rate is in agreement with the flow rate of geothermal water in active geothermal system ( 1 0 - 6 - t 0 .4 m/s) (Fujimoto, 1987). In this model, the area considered is (2 km x 0.5 km) and flow rate is 10 -4.2 m/s. When porosity is 2%, total mass flow rate is estimated as 10 42 m x 106 m 2 × 2/100 = 1.4 × 106 g/s. This is very similar to that in the Wairakei geothermal area in New Zealand which is 1.3 x 106 g/s (Elder, 1966).
1.4.6.6. Goldprecipitation due to mixing of fluids in epithermal system The mineralogy of hydrothermal alteration zoning, bulk compositional variation of altered rocks, thermodynamic consideration, mass transfer, and oxygen isotope computations, and the sulfur isotope study mentioned above all suggest that the mixing of two fluids (hydrothermal solution and acid groundwater) is the main cause for the geochemical and mineralogical variations in the Hishikari gold mine district. A large number of studies on the depositional mechanism of gold in epithermal system have been carried out (e.g., Shikazono, 1974a, 1986; Shikazono et al., 1990, 1993; Drummond, 1981; Reed and Spycher, 1985; Spycher and Reed, 1989). Depositional mechanism of gold in the Hishikari deposit has been discussed by several investigators (Izawa, 1988; Izawa et al., 1990; Shikazono and Nagayama, 1993; Nagayama, 1993b; Hayashi et al., 2000a,b). For example, Izawa (1988) and Izawa et al. (1990) thought that
200
Chapter 1
the mixing of ore fluids with groundwater, boiling and oxidation of ore fluids due to the interaction of ore fluids with oxidized hematite-rich paleosol are the main causes for the deposition and enrichment of gold in the veins. Hayashi et al. (2000a,b) suggested from a drastic change in 3180 of quartz at the precipitation sequence that electrum precipitated due to the mixing of fluids. The previous studies clearly demonstrated that Au thio complex is dominant among dissolved Au species in ore fluids responsible for Japanese epithermal A u - A g vein-type deposits (e.g., Shikazono, 1974a), considering the estimated f Q - p H range of Japanese epithermal gold deposits and the gold solubility due to thio complex (Seward, 1973, 1981). Thus, accepting the assumption that the Au thio complex is dominant among dissolved Au species in the ore fluids responsible for the Hishikari deposit, the precipitation reaction of gold in electrum is expressed as, Au(HS) 2 + H + + 1/2H2 ~ Au + 2H2S
(1-68)
This reaction suggests that a decrease in H2S concentration, and increases in H + concentration and fH2 (H2 fugacity) and temperature variations are important causes for the deposition of gold in electrum. It is also likely that the deposition of gold in electrum occurs by the following oxidation reaction. Au(HS) 2 + 1 5 / 2 0 2 + 1 / 2 H 2 0 --+ Au+2SO42 + 3 H +
(1-69)
This reaction suggests that oxidation of fluids is important as a depositional mechanism. However, this oxidation reaction seems difficult to explain the gold deposition from the following reasons. (1) The rate of oxidation of H2S to S O ] - is slow at a site of gold deposition. (2) O2 concentration of groundwater is very low. (3) HzS concentration in epithermal ore fluids seems higher than SO 2-. It is also possible that the following reaction is important for the deposition of gold. Au(HS)2 + 8 H20 ~ Au + 2 SO,]- + 3 H + + 15/2 H2
(1-70)
This reaction proceeds due to the degassing of H2, decreasing of H + and probably decreasing of temperature. We cannot evaluate this reaction as an important cause for gold deposition because no study on this reaction has been done. The importance of reaction (1-68) has been already cited by Reed and Spycher (1985) and Shikazono and Nagayama (1993). Shikazono and Nagayama (1993) favored the two fluids mixing as a cause for the gold deposition by the reaction (1-68) from the presence of acidic alteration zone overlying the Hishikari gold-quartz veins. For example, alunite and silicified rocks occur in higher elevations in the north-eastern part of the Hishikari district (Izawa et al., 1990). Shikazono (1985a) suggested based on the studies of wall rock alterations associated with Japanese epithermal-type Au-Ag deposits that acid sulfate fluids coexist with near-neutral chloride-rich fluids at relatively shallow zone from the surface (less than 1 kin) at the time of epithermal-type A u - A g ore formation. Reed and Spycher (1985) made a computation on the gold deposition from the mixed fluids. Their results support that gold deposition occurred in the Hishikari veins
Miocene-Pliocene Hydrothermal Ore Deposits
201
due to the mixing of two fluids (hydrothermal solution and acidic sulfate groundwater). However, it is uncertain that the alteration minerals (K-feldspar, K-mica and kaolinite) formed in the same stage of hydrothermal alteration. It may be possible that descending acid low-temperature solution (groundwater) formed kaolinite and cristobalite at different stage of gold-quartz mineralization associated with K-feldspar precipitation (Hedenquist et al., 1996). If such kind of mixing occurs, pH decreases, H2S concentration decreases, fH2 decreases and temperature decreases. These changes except the decrease of H2 proceed reaction (1-68). Another important depositional mechanism for gold in electrum is boiling of ore fluids, as inferred by many investigators (e.g., Drummond, 1981). Boiling of ore fluid causes an increase in pH. If dominant Ag species is AgCIf, this pH change as well as decreasing temperature, H2 degassing and dilution (decreasing of C1 concentration) causes the deposition of Ag in electrum according to the following reaction. AgC12 + H20 --+ Ag + 2 C1- + H + + 1/2 H2
( 1-71)
Galena, sphalerite and chalcopyrite precipitate also due to increase of pH due to the breakdown of base metal chloro complexes accompanied by the boiling by the reactions such as, PbC12 + H2S --+ PbS + 2C1- + 2H +
(1-72)
ZnC12 + H2S --+ ZnS + 2C1- ÷ 2H +
(1-73)
CuC12 + 2H2S + FeC12 --+ CuFeS2 + 4C1 + 4H +
(1-74)
However, in order to clarify the depositional mechanism of electrum and sulfides, more detailed description of alteration minerals, 3180, 3D data and the salinity (C1concentration)-enthalpy relationship are clearly required, and the two fluids mixing model has to be evaluated based on these data. It is difficult to evaluate that the boiling is an important mechanism for depositions of sulfides and electrum because of scarce data on fluid inclusions (salinity, temperature, enthalpy, 3180 and ~D). However, the above consideration and previous studies on epithermal Au-Ag deposits stressed an importance of boiling for gold depositions (Seward, 1991; Hayashi et al., 2000b). Therefore, it cannot be ruled out that the boiling and H2S loss are main causes for the depositions of electrum and sulfides (spbalerite, galena, and chalcopyrite) in the Hishikari deposit. However, preliminary study on the fluid inclusions indicates that boiling zone and ore zone containing high gold content are different (Etoh et al., 2001). Probably, the mixing of boiled fluid with acid groundwater caused efficient deposition of electrum in the Hishikari hydrothermal system.
1.5. Evolution of tectonics and hydrothermal system associated with epithermal and Kuroko mineralizations Numerous studies on the geologic and tectonic evolution in and around the Japanese Islands from Miocene to present have been carried out (e.g., Kitamura, 1959;
Chapter 1
202
Ozawa, t963; Sugimura and Uyeda, 1973). For example, Sugimura and Uyeda (1973) summarized volcanic rocks, degree of deformation of sediments, structural trends of sediments deposited, and amount of uplift and subsidence since Miocene. However, in contrast to these geologic and tectonic studies, very few studies on the relationship between tectonics and hydrothermal system in Neogene age have been carried out. Therefore, these studies are briefly summarized and then the relationship between geologic and tectonic evolution and evolution of hydrothermal system associated with the mineralizations (Kuroko deposits, epithermal veins) are considered below. 1.5.1. P a l e o g e o g r a p h y and stress field
Geologic environments and tectonic evolution of Green tuff region in Northeast Honshu have been studied by many workers (e.g., Kitamura, 1959). Paleogeography of the Japanese Islands and surrounding areas during a period of middle Miocene age (Kuroko-stage) and late Miocene-Pliocene age (vein-type stage) have been reconstructed in Fig. 1.145, mainly based on the paleontological studies (e.g., Ikebe, 1973, 1978; Chinzei, 1991). According to these studies, Green tuff region of middle Miocene age has been submarine environment in most areas (Fig. 1.145). Rapid subsidence occurred at middle Miocene age, which was contemporaneous with the age of Kuroko mineralization (Yamaji, 1990). From about 5 4-2 Ma uplift of Green tuff region in Inner Honshu province, Northeast Japan took place and this region became subaerial (Sugi et al., 1983; Otsuki, 1989, 1990).
(a)
(b)
o
K T 1 o*
,,¢(~
; @ushO
,.
L.V
0
200km
Land area -, -,~KKfi;hu Y £;k Coastal line *,v ~ ' ~,' (present-day) "~""
~3 Land area I
0
I
O
200km
Coastal line (present-day)
Figure 1.145. (a) Middle Miocene (Kuroko-stage)(solid circle: Kurokodeposits, open circle: Au-Ag vein-type deposits). (b) Late Miocene-Pliocene (vein-stage)(open circle: Au Ag vein-typedeposits) (Shikazono, 1987b).
Miocene-Pliocene Hydrothermal Ore Deposits
203
It is worth noting that the age of uplift is nearly coincident with the age of epithermal vein-type deposits in Northeast Japan including the Chitose, Yatani, Takadama, and Nebazawa Au-Ag vein-type deposits and Hosokura Pb-Zn vein-type deposits. However, the ages of some large base-metal vein-type deposits in Northeast Honshu (Ani, Osarizawa and Nikko Cu deposits) are older than this uplift, suggesting that unconsolidated marine sediments containing interstitial seawater were distributed in the mine area at that time. 334S data on sulfides and sulfates and high salinity of fluid inclusions from these large base-metal vein-type deposits support the above argument. Unfortunately tectonic situations of the regions other than Northeast Honshu of Neogene age are not well understood. However, it seems evident that even in the regions other than Northeast Honshu epithermal Au-Ag vein-type deposits formed when the uplift started and the area of land expanded. In addition to the paleontologic data, the country rocks of epithermal Au-Ag mine districts also suggest that epithermal Au-Ag vein-type deposits have formed under the subaerial condition: welded tuff occasionally occurs in the mine area (e.g., Sado, Nebazawa, Northeast Hokkaido) and in general submarine sedimentary rocks and volcanic rocks are poor or absent in the Au-Ag mine districts (e.g., epithermal Au-Ag vein-type deposits in Kyushu). As already noted, most epithermal Au-Ag vein-type deposits are hosted by young (late Miocene-Pliocene) volcanic rocks and by sedimentary rocks, but dominant host and country rocks for base-metal vein-type deposits are submarine sedimentary and volcanic rocks. Submarine felsic tuft', tuff breccia, dacite lava, intrusive rocks and mudstone are dominant host and country rocks of Kuroko deposits. Detailed studies on the change of stress field during Miocene to present in Northeast Japan have been carried out based on (1) the measurements of direction of dike (Nakamura, 1977; Kobayashi, 1979; Takeuchi, 1980, 1981, 1987), and (2) measurements of direction and age of epithermal vein-type deposits (Horikoshi, 1975b; Otsuki, 1989). Otsuki (1989) recognized the Neogene tectonic stress fields of the Northeast Honshu Arc by analyzing the epithermal veins and showed that in middle to late Miocene, ENE-trending al and a2 coexisted and a3 had a NNW trend, and it was replaced by E - W compression at 7 Ma in Southwest Hokkaido and at 5 Ma in Northeast Honshu. This view is supported by the analysis of trend of dikes (Tsunakawa and Takeuchi, 1986; Takeuchi, 1987) which indicates that the stress field changed from extensional to compressional during 8-6 Ma in southern Northeast Honshu (Figs. 1.146 and 1.147). These analyses of stress field suggest that the dip of subduction of Pacific plate during Miocene might have been steeper than during Plio-Pleistocene considering the mode of subduction of Pacific plate in Northeast Honshu (Niitsuma, 1979; Niitsuma and Akiba, 1984). Usually active plate subduction with gentle dip causes uplift of land and expansion of land area (Uyeda and Kanamori, 1979). Such change from steep subduction to a gentle one caused the changes in site of hydrothermal activity from submarine area during Miocene (Kuroko mineralization) to subaerial area by Plio-Pleistocene to present (epithermal vein-type mineralization). It is also noteworthy that major igneous and hydrothermal activities in the Japanese areas seems likely to have taken place almost at same times as the stress changes. The stress changes occurred at about 22 Ma, 15 Ma, 12 Ma and 8 Ma in the southern part of
Chapter I
204
3)
2)
...........
;~...... ..................... L .....
~
3
Ma
22~18
16-12 N
(I) -- ~fl
i' ' .;! ~
p , 6~(.~i.;....,,!S T11
0I
200km 1 TRENDOF DIKES NORMALFAULT •'~ REVERSEFAULT STRIKESLIPFLT. ~.','~,~STRESSTRAJECT.
12-8
6-(0)
Figure 1.146. Stress trajectory maps of southern Northeast Honshu in the late Cenozoic period, after Tsunakawa and Takeuchi (1986) with a slight addition. C~H...... trajectory is drawn by smoothing the inferred stress orientations from the selected dike-swarms with K Ar dates. Selected major faults with age estimation are also shown for indicating types of stress fields. T: Extensional stress field, where O'v > O-Hm,x >> OHma., and normal or gravity faulting is preferable. P: CompressionaI. O-H,,,~,~>2> CrHm,,~> OV, reverse or thrust faulting (Takeuchi, 1987).
Northeast Japan and at about 15 Ma in the eastern part of Southwest Japan (Tsunakawa, 1986; Takeuchi, 1987) (Fig. I. 146). The changes in stress fields, and intensities of igneous and hydrothermal activities seem to correlate to oscillatory motion of the Pacific plate (Jackson's episodes) (Jackson et al., 1975; Jackson and Shaw, 1975) (Masuda, 1984). Masuda (1984) and Takeuchi (1987) pointed out that the oscillatory motion of Pacific plate during the least 42 Ma correlates with magmatism, the intensity of tectonism, the change of stress field and the history of sedimentary basin in arc-trench system (Fig. 1.147). The above arguments also suggest that the mineralizations in arc and back-arc systems relate to the oscillatory motion of the Pacific plate.
1.5.2. Volcanic activity The type of volcanic activity in and around the Japanese Islands changed throughout Tertiary. In early Tertiary subaerial andesitic activity was intense. For example, in
Miocene-Pliocene Hydrothermal Ore Deposits AGE
EO.
Ma BP
40
OLIGOCENE
EARLYMIOCENE MIDDLE MIOCENE LATEMIO. PLIOCENEiQUAT.
30
Hltl
20
I"
15
'I'"
I
I
cSua
10
5
tl /1"'
"
]
ci
~
205
I
1.0 . 5
I
II/ll II
c2
d
~ AVE,AoErH
NO
OF HAWAIIANCHAIN
.~ ,-,
90t
cou,to,~lO~,,i,~
Episodes
0
o [ I
Inner Stress Field
r Hmax ORIENTATION
T
n .
"
(~n(P)
r .
"
P transverse
~
i
]
Figure 1.147. Jackson's curve and arc stress reorientations. Apparent swing motion of Pacific Plate (Jackson et al., 1975) and regional stress orientation at the Northeast Honshu convergent margin are illustrated in order to show their synchronous relationship. Dashed line represents the average trend of the Hawaiian volcanic chain. Pacific plate moves along the direction with fluctuation in reference to Hawaii Hot Spot. Vertically shaded parts of the graph indicate the climax phases of "clockwise episodes". Lower part of the figure shows the phases and reversals in orientation of tectonic stress fields on the inner zone of Northeast Honshu Arc (Takeuchi, 1987). Northeast Honshu, subaerial andesitic volcanic activity was dominant at Daijima stage (mostly early Miocene). F r o m middle Miocene b i m o d a l basaltic-dacitic activity started with rapid subsidence in Northeast Japan. The production o f volcanic activity was probably greater than today. Basalt at middle Miocene age erupted at Northeast Japan was studied by Shuto (1989) and Tsuchiya (1988, 1989) who showed that basaltic m a g m a generated in deep mantle. Dudfis et al. (1983) showed that p r e - K u r o k o ore basalt in the Hokuroku district has Mg number ( M g O / ( M g O + FeO*)), where FeO* is total iron content expressed as FeO, in the range 0.85-0.67 4-0.01, suggesting that the basalt is a relatively primitive, unfractionated, and mantle-derived melt. From late Miocene to present, subaerial a r c - v o l c a n i c activity ( c a l c - a l k a l i rocks, andesite, tholeiitic and high alumina basalt) started associated with uplift of the Japanese Islands. This volcanic activity is different from that at middle Miocene age. The above-mentioned changes in paleogeography, volcanism, crustal movement (subsidence and uplift), and stress field clearly demonstrate that these features of back-arc volcanism in e a r l y - m i d d l e Miocene are quite different from those o f Island arc volcanism in late Miocene to present. According to Yoshida and Yamada (2001), the age of change in volcanism from back-arc type to Island arc type in Northeast Honshu was 12.7 M a and this age corresponds to the age of Kuroko formation.
206
Chapter I
1.5.3. Tectonic influence on temporal and spatial relationships in Kuroko and vein-type deposits in southern Hokkaido, Japan It is worth studying Kuroko and vein-type deposits occurring in one metallogenic province. Shikazono and Shimizu (1993) carried out integrated geological, geochemical and mineralogical studies on Kuroko and vein-type deposits in a southwest metallogenic province which is described below. Geology of the province is composed of Paleozoic basements, Tertiary altered submarine volcanic and sedimentary rocks (Green tuff) and Quaternary volcanic rocks. The basements are shale, tuff, limestone and chert of unknown ages. A simplified geologic map is shown in Fig. 1.148. Tertiary rocks are distributed widely. They are composed of alternations of sandstone, mudstone, andesitic and dacitic tuff, tuff breccia and lava. These rocks are intensively and extensively altered and are called as Green tuff. Tertiary volcanic rocks are variable in composition. Andesite, dacite and basalt are found. Quaternary volcanic rocks are dominantly andesite lava and are abundantly distributed in the northern part of the province (Fig. 1.148). The vein-type deposits can be divided into two based on the metals produced; precious (Au, Ag) and base metal (Pb, Zn, Ag, Mn, Cu, Fe) vein-types. There are two sub-types of the base metal vein-type deposits, the C u - P b - Z n sub-type and the P b - Z n - M n - A g sub-type. C u - P b - Z n veins occur in southern part of the province. Large P b - Z n - M n - A g veins and A u - A g veins are distributed in northeastern part. In the northeastern part, A u - A g vein-type deposits occur in marginal zones of the province, while the base metal-rich deposits ( P b - Z n - M n veins and Kuroko deposits) in central zone (Fig. 1.149). The marginal zone is characterized by exposure of Quaternary volcanic rocks and Plio-Pleistocene volcanic rocks in which Au-Ag veins occur, whereas the central zone is by thick submarine volcanic rocks (Fig. 1.150), in which base metal-rich deposits (base metal veins and Kuroko deposits) occur (Fig. 1.150). Tertiary volcanic rocks, Quaternary volcanic rocks and faults are distributed, trending generally from NW to SE. Some C u - P b - Z n veins in southern part are hosted by basement rocks. On the other hand, P b - Z n - M n - A g and Au-Ag veins occur in Tertiary and Quaternary volcanic rocks. Recently K - A t age dating of ore deposits, associated volcanic rocks and plutonic rocks, have been carried out. These data are summarized in Table 1.25. It is obvious that the Kuroko deposits have been formed at middle Miocene, being very similar to the ages of the Kuroko mineralization at Hokuroku district, northeast Honshu where large Kuroko deposits occur. On the other hand, the vein-type deposits were formed during Plio-Pleistocene ages. It is also worthwhile to note that the ages and distribution pattern of the vein-type mineralization in this province are very similar to those of andesite which overlies the vein-type deposits (Watanabe, 1990b; Sawai et al., 1992b). Although ages of volcanic rocks which host the Kuroko deposits have not been determined, it is obvious that the Kuroko mineralization took place at the time of formation of volcanic rocks (mainly dacite). Therefore, it is clear that felsic volcanic activities were related to the mineralizations of the Kuroko deposits, while andesitic volcanism to the vein-type deposits.
Miocene-Pliocene Hydrothermal Ore Deposits 140°E
I
207
141,E
N
Inakuraishi
l
Akaiwa vein
ocene ks
Imai-lshizoki ~km
Basemen[ Rocks
Figure 1.148. Simplified geologic map and distribution of the Kuroko-typeand vein-type deposits in south Hokkaido (Shikazono and Shimizu, 1993). Open circle: precious vein-type deposits. Solid circle: base metal vein-type deposits. Solid square: Kurokodeposits. Minor elements associated with the vein-type and Kuroko deposits are different. Characteristic minor elements concentrated to the ore deposits are Se, Te, Hg, As, Sb and Bi in the A u - A g vein-type deposits, Ag, Bi, As, Sb, Sn, W and Mo in the base metal vein-type deposits, and Au, Ag, Sb, As, Mo and Bi in the Kuroko deposits. This difference in minor elements is consistent with that found in the other epithermal vein-type deposits in Japan (Shikazono and Shimizu, 1992). Analytical results of sulfur isotope previously obtained are summarized in Fig. 1.151.
Chapter 1
208
°; \ o ,"o o Q "
\
\.
/
10oB,o"~,~ s A P P o R o
6 Hg
~
ell O"a~\-
Sn
AgSn"13
~%'Mo \
TeSe~~ °
/~~/
\.
k
',Bi \.
,4"
o°
o.o
• CuPbZnMnVeins~ O AuAgVeins • Kurokoor massivebaritedep.
•
/:%
•
0
18
'\
1119
J
\
\/
2"0 , . - " 1
\
Figure 1.149. Distribution of ore deposits in northeastern part of the province (modified after Yajima, 1979). 1: Suttsu, 2: Kutosan, 3: Toyoha, 4: Akaiwa, 5: Matsukura, 6: Meiji, 7: Todoroki, 8: Teine, 9: Kobetsuzawa, 10: Otoyo, 11: Inatoyo, 12: Toyohiro, 13: Jozankei, 14: Toyotomi, 15:Koryu, 16: Eniwa; 17: Chitose, 18: Shiraoi, 19: Morono, 20: Minami-Shiraoi (Shikazonoand Shimizu, 1993).
It is found that ~348 values for different types of ore deposits are different; 4-1.8%o to 4-5.1%o for the A u - A g vein-type deposits, -2.8%0 to +7.5%o for the base metal veintype deposits and +4%o to 4-5%o for Kuroko deposits. 334S of the base metal vein-type deposits is widely variable, but generally higher than that of the A u - A g vein-type and nearly similar to that of Kuroko deposits. It is worthwhile to note that average ~34S values for the base metal vein-type deposits (except for the deposits in basements), A u - A g metal vein-type and Kuroko deposits in the province are identical to those previously obtained for the other metallogenic provinces in Japan (Shikazono, 1987b). It is also found that 334S values of the ore deposits are related to the host rock types; 334S of ore deposits hosted by basement rocks are relatively low. For example, 334S of the Imaiishizaki and Sasayama is -2.1%o and -0.9%0, respectively. 334S values of relatively large ore deposits such as the Toyoha, Jokoku, and Ohe are high in a range of 4-4%~ to 4-8%0, compared with those of small ore deposits. It is likely that sulfide sulfur in ore fluids responsible for small ore deposits is influenced by the surrounding rocks having low 334S values.
Miocene-Pliocene Hydrothermal Ore Deposits
209 141 °
140 ° 43 °
i~.! ¥ Vv
Quaternary
Pliocene
Miocene
Pre-Tertiary
subaerial volcaniclastic deposits subaerial lava & volcaniclastic rocks coarse laminated sandstone, subaerial to submarine lava & volcaniclastic rocks (Setana E) massive mudstone, submarine lava & volcaniclastic rocks (Kuromatsunai F.) mudstone & shale, submarine lava & volcaniclastic rocks (Yakumo E) granitic rocks sandstone, submarine lavas & volcaniclastic rocks (Kunnui E) mudstone & non marine sediments (Yoshioka F.) subaerial lava & volcaniclastic rocks (Fukuyama F.) ~ basement . . . . . . . rocks
Figure 1.150. Simplified geologic map of northeastern part of the Hokkaido metatlogenic province (modified after Yamagishi, 1989) (Shikazono and Shimizu, 1993).
Available homogenization temperatures of fluid inclusion from the base metal vein-type, A u - A g vein-type, and Kuroko deposits are summarized in Fig. 1.152. Salinity (NaC1 equivalent concentration) of inclusion fluids is 1-6 wt%, 1-14.5 wt% and 0 - 3 wt% for Kuroko deposits, base metal vein-type deposits, and A u - A g vein-type deposits, respectively. These data clearly demonstrate that the salinity of inclusion fluids for the base metal-rich deposits (base metal vein-type deposits, Kuroko deposits) is higher than that of the A u - A g vein-type deposits, while homogenization temperatures of fluid inclusion for these three types of ore deposits do not show a wide
Chapter I
210
Precious veins (/3
_>, c-< o .13
E Z
-5 -4 -3 -2 -1 0
1 2 3 4
5 6
7 8 9 10
(~ 34 8 ( % 0 )
B a s e metal veins 03 0'1 t~ c.< 0 ..Q
E Z
, l;>4Xb
t
-5-4-3-2-10
1 234
5 678910
1~ 34 8 ( % 0 )
Kuroko >09, c--
< "5
..Q
E Z
,
1
L
i
-5-4-3-2-10
~
i
1 2 3 4
56
7 8
910
1~ 34 8 ( % o )
Figure 1.151. Sulfur isotopic compositions of sulfides in the vein-type and Kuroko deposits. Solid box represents sulfur isotopic data from the ore deposits occurring in basement rocks (Shikazono and Shimizu, 1993).
Miocene-Pliocene Hydrothermal Ore Deposits
211
TABLE 1.25 K-Ar age data on the vein-type and Kuroko-type deposits in southern Hokkaido (Shikazono and Shimizu, 1993) Name of mine Vein-type deposits Chitose Koryu Toyoha Ohe lnakuraishi Todoroki Teine Ofukeshi Yakumo Jokoku Shizukari
K-Ar age (Ma) 4.7 3.6 3.5, 3.4, 3.3 1.1 2.2 3.3 3.3, 3.4 4.8, 4.9, 2.8, 2.7 2.1, 3.1 2.9 4.0 2.3 2.3 1.3 5.2 2.4
Name of mine Kuroko-typedeposits Kunitomi Horobetsu Kagenosawa Toya-Takarada Minamishiraoi Minamishiraoi
K-Ar age (Ma) 12.6 12.3 14.2 14.0 13.6 13.6
Disseminated-type deposits Hakuryu 6.5 Date 5.2
variation. Homogenization temperatures for the precious vein-type, base metal vein-type and Kuroko deposits are 180-280°C, 200-250°C, and 180-250°C, respectively. Ishiyama et al. (1987) estimated hydrogen and oxygen isotopic compositions of ore fluid responsible for the base metal vein-type deposits in Jokoku-Katsuraoka area, southwestern part of the province. Estimated 3180 and 3D values range from -49%o to -88%o and from -11.6%o to +5.5%o, respectively. Hattori and Sakai (1979) analyzed inclusion fluids from the Chitose A u - A g vein-type deposits and indicated that 3D and 31SO for the ore fluids are -65%o to -75%o and -4.8%o to -6.2%o, respectively. Interpretation of the geological and geochemical characteristics of the three types of ore deposits are given below. It is inferred that in the northern part of the province submarine volcanic rocks are thick in the central zone, while at marginal zone it is thin and the Plio-Pteistocene subaerial volcanic rocks are exposed. The vein-type deposits occur widely in the province. The precious vein-type deposits occur in relatively young (Plio-Pleistocene) volcanic rocks, while large base metal vein-type deposits (e.g., Toyoha, Inakuraishi, Ohe) and Kuroko deposits (e.g., Kunitomi) occur in central zone where thick Miocene submarine volcanic rocks are distributed (Figs. 1.149 and 1.150). Small base metal vein-type deposits occur in Paleozoic rocks in the southern part. 3348 of the vein-type deposits hosted by sedimentary rocks of the basement are low (less than 0%o), reflecting low 3348 of country rocks. However, they are scarce in number. 334S of the base metal vein-type deposits and Kuroko deposits are relatively high (average value; +4%0 to +5%o) and most probably influenced by sulfate sulfur of seawater
Chapter 1
212
15
d- 10 0 Z .=_
5
O9
0
2OO
30O
Filling Temp. (~C) Figure 1.152. Salinity (NaCl eq.wt.%) and filling temperatures for the base metal vein-type (solid circle), Kuroko-type (solid square) and precious vein-type deposits (open circle) in southwest Hokkaido (Shikazono and Shimizu, 1993).
trapped in the Green tuff rocks. ~34S of the precious metal vein-type deposits closely associated with Plio-Pleistocene subaerial volcanic rocks have igneous value (ca. +2%o), suggesting that sulfide sulfur was extracted from subaeriat country rocks. The causes for the different site of hydrothermal activity (submarine and subaerial environment) could be considered in terms of tectonic and geologic evolution of this metallogenic province from middle Miocene to Pleistocene. During Miocene age most of this province was in a submarine environment. Violent submarine volcanism (bimodal and basic type) took place at Miocene age in this province. This geologic environment may be related to an extensional stress regime (Uyeda and Kanamori, 1979). The Kuroko deposits have been formed related to this tectonic situation. From late Miocene, uplift took place due to the collision of Pacific plate to North American plate under the Kuril arc and Japanese island arc. Watanabe (1986, 1989, 1990a,b, 1991) studied the vein pattern, the age of veintype deposits and the volcanic rocks in southwest Hokkaido and showed that the major veins such as those at the Toyoha and Chitose have been formed at dextral strike-slip movement of an E - W trend, and those veins are situated at the west-southwest extension of the maximum displaced zone within the dextral shear belt along the Kuril arc. Watanabe (1990b) also showed that the veins in the Sapporo-Iwanai district strike E - W and are oblique to the N W - S E volcanic chains which are sub-parallel to the maximum principal stress estimated in southwest Hokkaido during Late Miocene to Holocene and oblique subduction of Pacific Plate was active during the Plio-Pleistocene age. Yamagishi and Watanabe (1986) studied the geologic faults, the dykes, the veintype deposits and active faults in the rocks of the middle Miocene to Quaternary in
Miocene-Pliocene Hydrothermal Ore Deposits
213
this province and clarified that in the middle to late Miocene, southwest Hokkaido was characterized by a tensional stress field of an E - W to N W - S E direction, while in the late Miocene to Quaternary this area was in a compressional stress field. They showed that the main ore veins are recognized as sheared fractures arranged mostly in an E - W direction indicating the signal trajectory lines of an E - W to NW direction. It is noteworthy that the stress field of N W - S E direction estimated from a province of dikes during Miocene is parallel to the distribution trend of the Kuroko deposits in northeastern part of the district. Yamagishi and Watanabe (1986) suggest that the dip of subduction of Pacific plate during Miocene might have been steeper than during Plio-Pleistocene, considering the mode of subduction of Pacific plate in Northwest Honshu (Niitsuma, 1979). Usually active plate subduction with gentle dip causes uplift of land and expansion of land area (Uyeda and Kanamori, 1979). Such change from steep subduction to gentle one caused the changes in site of hydrothermal activity from submarine area during Miocene to subaerial area by Plio-Pleistocene to present. This change from extensional stress regime to compressional regime may correspond to the change of mode of subduction from Mariana-type to Chilean-type as defined by Uyeda and Kanamori (1979). It has been pointed out that this change has occurred at about 5 Ma in Northeast Honshu (Sugi et al. 1983; Ohmoto et al., 1983), and about 7 Ma in southwest Hokkaido (Otsuki, 1989). 1.5.4. Geochemical features of sedimentary rocks formed in the Japan Sea as a proxy for hydrothermal activity As noted already, intense submarine hydrothermal activity took place in the Japan Sea in 15-12 Ma, associated with Kuroko mineralization. However, it is uncertain that submarine hydrothermal activities associated with the Kuroko mineralization took place in the other periods from middle Miocene to present in the Japan Sea. Therefore, the geochemical features of sedimentary rocks which formed from the Japan Sea at these ages have been studied by the author because they are better indicator of age of hydrothermal activities than those of hydrothermally altered igneous rocks because the samples of continuous age of sedimentation are able to be collected and the ages are precisely determined based on microfossil data (foraminiferal, radioralian and diatom assemblages). Thick sedimentary pile from middle Miocene to late Pliocene is exposed in the Oga Peninsula, northern Honshu, Japan (Fig. 1.153). Age of the sedimentary rocks has been determined by microfossil data. Thus, the sedimentary rocks in the Oga Peninsula where type localities of Miocene sedimentary rocks in northern Japan are well exposed have been studied to elucidate the paleoenvironmental change of the Japan Sea (Watanabe et al., 1994a,b). Kimura (1998) obtained geochemical features of these rocks (isotopic and chemical compositions) and found that regional tectonics (uplift of Himalayan and Tibetan region) affect paleo-oceanic environment (oxidation-reduction condition, biogenic productivity). However, in their studies, no detailed discussions on the causes for the intensity and periodicity of hydrothermal activity, and temporal relationship between hydrothermal activity, volcanism and tectonics in the Japan Sea area were discussed. They considered only the time range from ca. 14 Ma to ca. 5 Ma.
214
Chapter 1
NYUDOZAKI
SARUgAWA
i
NISHIKUROSAWA ~
NOMI/RA
K1TAURA
ANDEN L"
' RECLAIMED
LAND
AIKAWA
TOGABAY~_I~j ~NINIOME.G ..A. .TAICHLNOMEG . ATA S A N ~ O M E -GATA
KAMO
Mr, K E N A S H I Y A M A JAPAN SEA
MONZEN
SUGOROKU
MINAMIH IRt..qAWA 0NNAGAWA
5kin
['~
~
Kampu-zanVOlCaNOs
Megata volcanic ejecta
~ KatanishiFormation ~ ~'/tJ Togapumicebed i ~rl~ ShibikawaFormation ~ ~
FunakawaFormation OnnagawaFormation NishikurosawaFormation DaijlmaFormation
WakimotoFormation ~
MonzenFormation
Kitaura Formation
AkashimaFomation
~
Figure 1.153. Geologic sketch map of the Oga peninsula.
Therefore, the wider time range from middle Miocene to present is considered below based on available age data on hydrothermal ore deposits (Kuroko deposits, epithermal vein deposits) and hydrothermal alteration in the mine areas in Northeast Japan.
Miocene-Pliocene Hydrothermal Ore Deposits
215
The geochemical features of the sedimentary rocks in the Oga Peninsula and the hydrothermal activity in Japan Sea deduced from these features are described below. The area is located at the west of Oga Peninsula (Fig. 1.153) and is composed of Miocene sedimentary rocks (Nishikurosawa, Onnagawa and Funakawa Formations). The Nishikurosawa Formation is composed of siltstone, mudstone, conglomerate and sandstone. Siltstone and mudstone contain foraminiferal fossil such as GIoborotalia birnageae, and G. denseconnexa, indicating Zone N. 9 by Blow (1969). The upper part is characterized by glauconite-bearing sedimentary rock. The total thickness is about 150 m. The Onnagawa Formation conformably overlies the Nishikurosawa Formation and is composed of siliceous shale and shale. The rocks are characterized by organicrich laminated diatomaceous deposits, siliceous microfossils, and fish bones, while foraminiferal fossil is poor in amounts. The total thickness is about 300 m. The age of base of the Onnagawa Formation is estimated to be 12.9 Ma based on diatoms (Koizumi and Matoba, 1989). The age of the top of the Formation is 5.8 Ma. The Funakawa Formation conformably overlies the Onnagawa Formation. The thickness is about 1000 m. The Formation is composed of siltstone, intercalated by tuff and tuffaceous siltstone. Foraminiferat fossil and siliceous microfossil exist in the formation. Average analytical data on the sedimentary rocks are compared with NASC (North American Shale Composite; Gromet et al., 1984) in Fig. 1.154. Average chemical compositions of the Nishikurosawa are very similar to NASC. However, Fe203, MgO, TiO2, MnO and P205 contents are higher and K20 content is lower than those of NASC, respectively. SiO2 and A1203 contents are similar with each other. The Onnagawa Formation is characterized by very high SiO2 content (average 89.12 wt%). Most of the other major element contents are lower than those of NASC, although MnO and P205 contents of the Onnagawa are higher. Average major element contents except SiO2 content of the Funakawa are very similar to NASC, but SiO2, Na20 and MnO contents are higher and the other element contents are lower than the NASC. Analytical data on minor element contents are compared with Average Shale by Turekian (1972) (Fig. 1.155). The average contents of most of the elements except Zn and Ba of the Nishikurosawa are same in orders to NASC. Average contents of minor elements of the Onnagawa are mostly lower than Average Shale. However, Cu, Zn, Mo, Ba and U contents are anomalously high compared with Average Shale. Average contents of minor elements of the Funakawa are lower than NASC which is due to slightly higher SiO2 content of the Funakawa. REE pattern normalized by NASC (REE contents; Goldstein and Jacobsen, 1988) are shown in Fig. 1.156. Nishikurosawa Formation is characterized by Eu positive anomaly which is calculated by the following equation: Eu/Eu* = (2Eu/Eu~ASC)/(Sm/SmNASC -k-Gd/GdNASC)
(1-75)
Chapter 1
216 (wI%}
I00
~A)
N/shik~t,
shale
rasa~;,~
8 ~g 40 NK¢-~eries s v e r a z e
2{) l{
-
@rome~. e~
~L 1984]
t :
.---
10E
(B) O n n a g a w a harod s h a l e
......~
OO-ser es @r~met
\
N
80
4O ~ver~e
et d
.
i9~
\
\
-
\ \
I
2
Miocene-Pliocene Hydrothermal Ore Deposits
217
4~ 'i } t{
k
Figure 1.154. Major element composition of studied rock samples. (A) Nishikurosawa shale, (B) Onnagawa hard shale, (C) Funakawa shale. Shaded area represents ranges.
The REE pattern for the Funakawa plot within 0.1-1 of NASC normalized value (Fig. 1.156). Ce and Eu anomaly for the Funakawa are weak (Ce/Ce* = 1.20 4-0.01; Eu/Eu* = 1.09 4- 0.05) (Fig. 1.157). The REE pattern for the Nishikurosawa exhibits slightly light REE enriched one (La/Yb = 0.55 4- 0.09) and plots close to the line of NASC normalized value -----1. Positive Ce anomaly is also found in the Nishikurosawa. This anomaly is defined by the following equation. Ce/Ce* = (3 Ce/CeNAS¢)/(2La/LaNASC + Nd/NdNAsC)
(1-76)
Hydrothermal solution venting from midocean ridges and back-arc basins has positive Eu anomaly (Klinkhammer et al., 1983; Michard et al., 1983; Mitra, 1994; Shikazono, 1999a) (Fig. 1.158). Therefore, the positive Eu anomaly of the sedimentary rocks is thought to be due to a contribution of hydrothermal solution. In order to know the contribution of hydrothermal solution the positive Eu anomaly of seawater (Eu/Eus*awater) is useful.
Chapter 1
218
' ppm~ f1
C
i
]ktrcB~ 1~;2]
i
I
[
~y °
~0~
m a h a ~ d sAale
fB~ 0 7 ~
i0/
° b" ......
%#
{ J
V
~n
I
,¢
{ h
~r ~e~
Zr y
?~ Nb
A ..........
Hf
R
Ba
Ph %~
U
Miocene Pliocene Hydrothermal Ore Deposits
219
(ppm) }£
[ ........ r - - - 7 - - 7 - - - ]
.............................................................
~ z t ;k~ a~
"
~
...........
192°9.
i
[ I} ~--
~
ot
Figure 1.155. Trace element composition of studied rock samples. (A) Nishikurosawa shale, (B) Onnagawa hard shale, (C) Funakawa shale. Shaded area represents ranges. The anomaly at that time is estimated based on the following equation. gu/gUr*oc k = n (gu/gut%rrigeneous) -Jr (1 - n)(gu/gus%awater )
(1-77)
where n is proportion of terrigeneous component in sediments, Eu/Eur*ck is Eu positive anomaly of sedimentary rocks, and Eu/Eut*rrigeneous is positive Eu anomaly of terrigeneous component. Eu/Eus*awater estimated based on the above equation is shown in Fig. 1.159. It is evident in this figure that intense hydrothermal activity occurred in the Nishikurosawa stage (14-13 Ma), it was very high " (Eu/Euseawater * -- 2.4) at the Nishikurosawa/Onnagawa boundary (12.6 Ma) which is similar to that of the sediments below the hydrothermal pool of Red Sea (Fish debris, Eu/Eu* = 2.4-4.9; Oudin and Cocherie, 1988), the intensity of hydrothermal activity decreased in the Onnagawa stage towards the Funakawa stage, and intense hydrothermal activity occurred intermittently at ca. 12.6, ca. 10.5 and ca. 8.2 Ma. High Zn, Cu and Ba contents of the Nishikurosawa Formation also indicate that the intense hydrothermal activity occurred at these ages. Magnetic susceptibility data are inferred to have been reflected by hydrothermal activity. The magnetic susceptibility data on the sedimentary rocks are shown in
220
Chapter 1 10
Funakawa shale
0.1
0,01
l
i
l
i
l
l
l
l
l
l
l
l
l
l
10 Onnagawa hardshale N
o~
@
=, 0.1
r/3
.< Z
l
l
l
f
i
l
i
l
l
l
l
I
10
0,01
Nishikurosawa shale
O.I
(1,(11
I
La Ce
I
t
I
I
v
NdSmEuGd
,
~
Dy
I
i
Er
I
E
YbLu
Figure 1.156. NASC-normalizedREE patterns. REE data of NASC obtained by Goldsteinand Jacobsen (1988) are used as normalizationvalue.
Fig. 1.160. The magnetic susceptibility values of the Nishikurosawa sedimentary rocks are high but those of the Onnagawa are relatively low. The values of the Funakawa are lower than those of the Nishikurosawa but higher than those of the Onnagawa. The values around the Nishikurosawa/Onnagawa boundary are the highest and they are variable in the Nishikurosawa and in the Onnagawa. The values of the Onnagawa are relatively high at 14-13 Ma, 12.6 Ma, 10.5 Ma, and 8.2 Ma. This variation correlates to Eu/Eus*awater
Miocene-Pliocene Hydrothermal Ore Deposits 1.4 1.3
221
m
Eu/Eu, 1 J:o
o
ml
|E
o o
E
1.2
0
o ®
m B
o
0
e! a o
1.1
o
o no a n o m a l y ~
1.0
o
o
O e
Pl )sitive negative
. . . . . . . . . . . . . . . . . . . . .
0.9 NK-shale
0t-t3
OG-hard
! .... l iFK-shale 01-04
shale 01-21
Figure 1.157. Eu/Eu* values of studied rock samples. I
w -E
0.01
I
I
I
1
I
Z
I
~ 13*N
-~
o
J
J
1984
-
13°N F982
m 2 1 * N 1981 ~.
~ \
O
o,oo :
(xlO-3t
\
_
_
#.°
-
/\
J
\
E
oc-
-
/
Cl~
0 0.0001
-
-
-
i
t
Ce
Nd
I
I
L
Sm Eu Gd
L
I
I
Oy
Er
Yb
Figure 1.158. Rare earth elements in vent fluids from the East Pacific Rise (redrawn from Michard and Albarede, 1986). The Sm values for the 1984 I3°N set are thought to be somewhat too high because of contaminationfrom the samplingsyringes.Data for basalt from 13°N East Pacific Rise are from Michard et at. (1983) (Scott, 1997).
variation, suggesting that high magnetic susceptibility may be due to high content of magnetite in the sedimentary rocks. Magnetite is considered to have formed biogenetically because magnetite is very fine-grained. If the grain size of magnetite is less than 0.03 Ixm, magnetic susceptibility of sediments containing magnetite is very high (Ioka and Yamazaki, 1994) and magnetite is thought to be of biogenic origin (Yamazaki et al.,
Chapter 1
222 (A)
(B)
- 2.0 " ~ ~ [ ' ~ t ~
......., ~
H ydrothermaI sediment
Seawater
d~tal
* - l.OI--n..... 0.8 I
Modern*
]
[
'v~"~O.~Q~Q " "~O'
aly
negative >12.9
., ,
12
,
11
,
10
,
9
"
% ,
8
7
5.8>
Middle Miocene to early Pliocene
J~l~r'. (Ma)
]
Figure 1.159. Eu/Eu* values of (A) modern sediment, hydrothermal solution and seawater and (B) midMiocene to early Pliocene Japan Sea (gu/Eus*awater, see in text). Modern data are from the Pacific ocean, except fish debris in the hydrothermal sediment (Red Sea). Modern* is defined here as including Quaternary time. Data sources: Hydrothermal solution (Michard and Albarede, 1986), seawater, fish debris in the Atlantic II deep (Oudin and Cocherie, 1988), hydrothermal sediment (Ruhlin and Owen, 1986).
1991). Probably, iron of biogenic magnetite was originated from hydrothermal solution. It is considered that ferric iron of hydrothermal solution was oxidized by iron oxidizing bacteria to form magnetite. The ages of Neogene mineralization and hydrothermal alteration in and around the Northeast Honshu and Hokkaido have been determined by K - A r data on K-minerals (K-feldspar, sericite). These data are summarized in Fig. 1.147 and Table 1.26. It seems clear by comparing Fig. 1.159 with Table 1.26 that the ages of hydrothermal mineralization and alterations determined by K - A r age dating are consistent with those of sedimentary rocks affected by hydrothermal activity in the Oga. Hydrothermal activities were intense at ca. 14-13 Ma, 12.6 Ma, 10.5 Ma, and 8.2 Ma. As already noted, major igneous and hydrothermal activities in the Japanese Islands seems likely to have taken place nearly at same times as the stress changes (Fig. 1.147). The stress field changes occurred at about 22 Ma, 15 Ma, 12 Ma and 8 Ma in the southern part of Northwest Japan and at about 15 Ma in the eastern part of Southwest Japan (Takeuchi, 1987) (Fig. 1.147). Kuroko and vein-type mineralization occurred at 23 Ma (Sado), 15-14 Ma (Kuroko in Hokuroku, Sado), 12.8 Ma (Karuizawa), 8 Ma (Takadama), 5 Ma (Hosokura), 3 Ma (Yatani) (Shikazono, 1985e, 1987b; Shikazono and Tsunakawa, 1982; Sugaki et al., 1986; Sawai and Itaya, 1996; Otsuki, 1989).
Miocene-Pliocene Hydrothermal Ore Deposits
223
(10 -3 SI) m!
1.0 t
Magnetic
2
susceptibility 0.8 0.3
J
Q m
0.2 [] WW
0.I
0 ® 0
0
0
®® @ ®
0 i
t
t
I
i
I
I
I
t
NK-shale 01-13
i
;
I+
I ~79
I ®I [ I I I I I I l0I T I I I I
OG-hard shale 01-21
0 fin Oi @ II FK-shale 01-04
Figure 1.160. Magnetic susceptibility value of studied rock samples. Most younger two samples (NK-12 and -13) of Nishikurosawa shale have anomalously high magnetic susceptibility value.
The changes in stress fields, and intensities of igneous and hydrothermal activities except 12.6 Ma seem correlate to oscillatory motion of the Pacific plate (Jackson's episodes; Jackson and Shaw, 1975; Jackson et al., 1975) (Masuda, 1984). Masuda (1984) and Takeuchi (1987) pointed out that the oscillatory motion of Pacific plate during the last 42 Ma correlates with magmatism, the intensity of tectonism, the change in stress field and the history of sedimentary basin in arc-trench system (Fig. 1.147). The above argument also suggests that the hydrothermal mineralizations in arc and back arc systems relate to the oscillatory motion of the Pacific plate. It is generally accepted that Kuroko deposits formed under the submarine environment, while polymetallic vein-type deposits in central and Northwest Japan (Ashio, Tsugu, Kishu, Obira, etc.) under the subaeriat environment. This spatial difference in the distribution of back-arc and vein-type deposits is found also for present-day mineralizations. For example, back-arc deposits are forming at Okinawa Trough, while base-metal and precious-metal precipitations are occurring on land such as at Ibusuki and Beppu geothermal area, Kyushu. Recent gold mineralizations (ca. 1 Ma) occurred at Noya and Hishikari areas in southern Kyushu. These mineralizations can be regarded as almost contemporaneous mineralization with the Okinawa mineralization. This kind of temporal and spatial relationship between epithermal Au vein-type mineralization and back-arc mineralization are found also in the Izu-Bonin area. Seafloor
Chapter 1
224 TABLE 1.26
Isotopic ages of metalliferous veins, altered country rocks and the igneous rocks which have genetical relations to the mineralization (Otsuki, 1989) Mine
Age (Ma) Mineral
Sanru Ohe Toyoha Inatoyo-Toyohiro Chitose Horobetsu Kagenosawa Koryu Hakuryu Date Ani Koaizawa Kameyamamori Arakawaohsawa Rata Sugisawa Kakkonda Hosokura Yatani
Takatama Karuizawa Ohizumi Sado
Ashio Nehazawa Chichibu
12.4 3.3 2.2 4.7 3.6 12.3 14.2 1.1 6.5 5.2 11 12.6 8.1 13.4 9.4
5.8 3.3 3.4 3.6 8.4 12.8
Metals
Vein strike
Au.Ag Mn>>Zn-Pb Zn > Pb>>Mn
N70°E, N85W; N15W N55°-85°W N40°-50°W, EW NS N65°E-N80W
Rock 6.6 8.4 8.5 3.4 10.3
12-15 23.6 15.2 8.7 13.6 12.0 4.9 9.7 9 5.2
5.3 13.4 14.5 22.1 24.4 14.8 5.0-5.7
Au-Ag Au-Ag-Cu.Pb.Zn Au.Ag.Cu-Pb.Zn Au-Ag Au.Ag.Cu.Pb.Zn Au-Ag-Cu-Pb.Zn Cu Ag.Cu.Pb.Zn Au.Ag.Cu.Pb-Zn Zn.Pb Au.Ag.Cu.Pb-Zn Au-Ag-Cu-Pb.Zn
N70°E, N55°-80°W N70° 90°E N80°-90°E ENE ENE N23°-70°E, N80°W N50°-60°W N25°W WNW-EW-NS N45°E, EW, NS NW N30°-80°E NNW-NW
Zn.Pb Zn.Pb
ENE, EW, NW EW, NE, NW
Pb.Zn Au-Ag Au.Ag Au.Ag Ag.(Pb-Zn) Zn.Pb.Cu Au.Ag
EW, 70W N70.W-EW-N45E
Cu Au-Ag 6.6
WNW, ENE, NNW N60°W, N60°E EW-ENE + NS + NW N65°-90°E, NW
N50°-70°E, EW N70 °, 85°E NS
b a s e - m e t a l m i n e r a l i z a t i o n s i m i l a r to the K u r o k o m i n e r a l i z a t i o n is f o u n d in I z u - B o n i n s e a f l o o r ( l i z a s a et al., 1999), w h e r e a s e p i t h e r m a l g o l d v e i n - t y p e m i n e r a l i z a t i o n o c c u r r e d at v e r y r e c e n t age (ca. 1 M a ) at Izu P e n i n s u l a , c e n t r a l Japan. T h i s spatial d i f f e r e n c e is c o n s i s t e n t w i t h the d i s t r i b u t i o n w i t h that o f h y d r o t h e r m a l d e p o s i t s o f m i d d l e M i o c e n e ( K u r o k o a n d p o l y m e t a l l i c v e i n - t y p e d e p o s i t s in Japan). I f the s u b m a r i n e vs. s u b a e r i a l h y p o t h e s i s m e n t i o n e d a b o v e is correct, it is e x p e c t e d t h a t K u r o k o d e p o s i t s are f o u n d in the m a r i n e s e d i m e n t a r y a n d v o l c a n i c h o r i z o n s w h o s e
Miocene-Pliocene Hydrothermal Ore Deposits
225
ages are 12.6 Ma, 10.5 Ma, 8.5 Ma and 5 Ma as well as 15-14 Ma (Kuroko deposits in Hokuroku area) and present-day (Okinawa and Izu-Ogasawara deposits). Large Kuroko deposits formed at 15-14 Ma (ore deposits in Hokuroku and Hokkaido) and at present (Okinawa Trough, Izu-Ogasawara). These ages correspond to rapid change of oscillatory motion of Pacific plate (Jackson's episodes). The change in the motion of the Pacific plate occurred at 15-14 Ma, 10.5 Ma, 8 Ma, 6 Ma, 3-2 Ma and probably present (Fig. 1.161). Thus, small Kuroko deposits probably formed at submarine environment at 10.5 Ma, 8 Ma, 6 Ma and 3-2 Ma, although these ore deposits have not been discovered. However, it is noteworthy that the 12.6 Ma corresponding to the Nishikurosawa/Onnagawa boundary when intense igneous and hydrothermal activities occurred does not correlate to Jackson's episode. The geochemical characteristics of the Nishikurosawa/Onnagawa boundary (12.6 Ma) are distinct from those of the other ages affected by hydrothermal activity (15-14 Ma, 10.5 Ma, 8.5 Ma); The positive Eu anomaly and magnetic susceptility were the highest compared with the other ages, glauconite is abundant, paleoenvironment was extremely reducing, and mass extinction occurred. This difference suggests that the cause for the volcanism and hydrothermal activity at that time were different from the other ages. The cause for hydrothermal and igneous activities at 12.6 Ma is uncertain because of no relationship with Jackson's episode. However, it is speculated that bolide impact may affect igneous and hydrothermal activities at this age. In addition to no relationship with Jackson's episode, this speculation comes from that the Earth's surface environmental change at the other geologic boundaries such as Eocene/Oligocene which is considered to have been influenced by bolide impact (Keller, 1986; Montanari et al., 1993; Vohof et al., 2000). Bolide impact may be important. But, further work should be clearly necessary to evaluate this hypothesis which is highly speculative.
1.5.5. Mode of subduetion and formation of back-arc basin
Uyeda and Kanamori (1979) divided mode of subduction into two types: Marianatype characterized steep subduction and Chilean-type characterized by gentle subduction estimated from the dip of Benioff-Wadati zones (Fig. 1.161). The geological phenomena associated with these subductions are shown in Fig. 1.162. It is inferred that the change in mode of subduction from Mariana-type to Chilean-type occurred at ca. 5 Ma in Northeast Honshu. Kuroko deposits are associated with Mariana-type, whereas epithermal vein-type deposits (particularly Au-Ag deposits) with Chilean-type. The above-mentioned view on the evolutionary history in metallogenic provinces from the formation of massive sulfide deposits at early-stage to that of the vein-type deposits in late-stage is consistent with the view by Cathles (1986) who has shown that massive sulfide deposits have been formed at an earlier stage than gold-quartz veins in Archean greenstone belt. Therefore, it is likely that the change in style of mineralization from massive sulfide deposits to gold vein-type deposits caused by the change in mode of subduction is common geologic phenomenon from Archean to Quaternary age in an Island arc-back-arc geologic setting. However, the relationship between the mode of
Chapter i
226 . Andaman M. America (12 N), C. Aleution, L. Antitles 100 km -
"~--'~
~
S.Chile(40%) Sumatra(5~S)
/S.Chile (30"S)/
I~~__~.~ru
~,laska,~.
'
~
New
~ ~
~.
j
(10°S) N.Chile(20°S)
Hebridesd'~ ~k'~~.~ "% ~
0
Depth, km 100 200
.yo.yu.p'%\'<,
'New Zealand
Ik ~
~
• Solomons
~ ~It~ , ~
Philippine-- - - ~ N e w Britain
~ ~
300
'~ '~
NEJalpan 400 ~-,,~Kurl~,
Java"
Mariana
'
'
700
Figure1.161.DipofBenioff-Wadatizones(UyedaandKanamori,1979). subduction and style of mineralization during the geologic history is not clear and should be studied in future. Horikoshi (1975a) proposed the hypothesis that Kuroko deposits formed at the time changing from extensional environment to compressional environment. He thought that metal-bearing fluids in the crust were pushed out by the compression. In contrast, Otsuki (1990) and Yamaji (1990) thought that back-arc rift activity and rapid subsidence occurred at ca. 15 Ma under the extensional environment. Bimodal volcanism was dominant at that time. Therefore, it is inferred that Kuroko formation took place at extensional environment, rather than compressional environment favored by Horikoshi (1975a); formation of Kuroko deposits may be related to the opening of the Japan Sea. However, the relationship between the opening of the Japan Sea and Kuroko mineralization is unclear, and there are many opinions on the genesis of the Japan Sea and timing of opening of the Japan Sea. For instance, Otofuji et al. (1987) and Otofuji (1996) showed that Southwest Japan Sea opened at middle Miocene very rapidly and widely based on paleomagnetic data. However, the rate of opening of the Japan Sea estimated by Otofuji et al. (1987) is too high (60 cm/year) and seems to be unreasonable. Therefore, many workers think that opening of the Japan Sea did not occur very rapidly and widely at middle Miocene age, but it occurred from earlier age (25 Ma) and opening occurred several times by the slower rate (Nakajima and Nakagawa, 1994). Twice openings are favored by several workers (Nakajima and Nakagawa, 1994). It is considered that there were several axes of the opening based on K - A r ages of volcanic rocks (Kobayashi, 1983).
Miocene-Pliocene Hydrothermal Ore Deposits
Ce_,
¢
-,,,
227
d
.,o*o
Mar~aria
"~~~ "
Figure 1.i62. Two kinds of subduction boundaries (Uyedaand Kanamori, 1979).
In contrast to Southwest Japan, opening of northeastern part of the Japan Sea is unclear, compared with the southwestern part of Japan. Tosha and Hamano (1988) made a paleomagnetic study of Tertiary rocks of Oga Peninsula (northern Honshu) and considered that a counterclockwise rotation of Northeast Japan with respect to eastern Asia took place between about 22 Ma and 15 Ma and the before the rotation, Northeast Japan was situated along the east coast of the Asian continent. Lallemand and Jolivet (1986) have interpreted the opening of the Japan Sea as a pull-apart basin between two right-lateral strike-slip fault zones; the Yagsan-Tsushima fault to the west and the Tartary-Hidaka shear zone to the east. Southward migration of Southwest Japan as a drawer between a right-lateral system to the west and a left-lateral one to the east along the Tanakura Tectonic Line (Otsuki and Ehiro, 1978) or a clockwise rotation of Southwest Japan about a pole located in the Tsushima Strait (Otofuji et al., 1987), or a combination of both (Kobayashi, 1983) have been proposed.
228
Chapter 1
Although many discussions on the origin of the Japan Sea have been carried out since the pioneer works by Tokuda (1927), and Terada (1934) who considered that the Japan Sea formed by southward migration from the Asian continent, the problem of the mechanism of the opening of the Japan Sea remains unsolved as mentioned above. Kuroko formation occurred at middle Miocene (15-16 Ma). But the opening of Southwest and Northeast Japan Sea occurred probably from the age earlier than this age. Therefore, it is thought that the beginning of the opening of the Japan Sea did not directly relate to the Kuroko formation. Cathles (1983a) proposed failed rift hypothesis to explain unusual features of the Green tuff region and Kuroko deposits. He mentioned that this hypothesis can account for the distribution of mine districts in the Green tuff region, the observed extensive and substantial premineralization subsidence and postmineralization uplift in the region, and its volcanic evolution. He further pointed out that the Green tuff belt of Japan is strikingly similar in geology, tectonics, and mineral deposits to Archean greenstone belts, suggesting that Archean belts may be failed rifts. As mentioned above, formation of back-arc basins and marginal seas may be important for the formation of Kuroko and vein-type deposits, although genetic relationship between Kuroko formation and opening of the Japan Sea is not clear. For example, Horikoshi (1977) insists that vein-type deposits in Northeast Hokkaido did not form without the opening of Ohotsuku back-arc basin. In section 2.3 and in Chapter 3, it is shown that the formation of back-arc basins take important role for the mineralization (back-arc deposits (Kuroko deposits), epithermal Au veins) and global geochemical cycle. Thus, it must be worth considering the formation mechanism of back-arc basins. The origin of the back-arc basins has been investigated considerably (Karig, 1971; Sleep and Toks6z, 1971; Uyeda and Kanamori, 1979; Tamaki and Honza, 1991; Uyeda, 1991; Tamaki, 1995) and various explanations for the origin have been proposed. Tamaki and Honza (1991) summarized the previously proposed models and the currently plausible models of back-arc spreading (Fig. 1.163). Model 1 shown in Fig. 1.163 is a slab-induced upwelling model. Upwelling generated along the down going slab-mantle boundary (Karig, 1971) or by secondary convection that is introduced by the slab (Sleep and Toks6z, 1971) causes back-arc formation (Karig, 1971 ). Model 2 is a plume injection model. Miyashiro (1986) showed that a northward migration of "hot region" (a large high-temperature region at a deeper level) governed the formation of back-arc basins in the Western Pacific. Tatsumi et al. (1990) examined the distributions of Cenozoic basalts and active rift systems in the Northeast China region and Quaternary subduction unrelated volcanism in the northwestern part of New Zealand (Fig. 1.164) and concluded the following: (1) Injection of the asthenosphere into the mantle wedge and the dam-up effect of the subducted slab explains the rifting process in the Japan Sea (Fig. 1.165). (2) Injection of the asthenosphere was associated with the eastward horizontal convective flow caused by the upwelling of asthenosphere beneath the northeastern China region (Fig. 1.165).
Miocene-PIiocene Hydrothermal Ore Deposits
229 , J / ~ - ""--I
I=IS!~!I, ! i I ~ k ~ I m tb mOdi~ £ 8 (le
Mode 1i Ssb4ii :d c@d pw@t gmsd~i!
XXXX
X
XX
M~de 4. ILIt,q!! £ iO@l£
m~d~l ? I~+I
M s I e 2 P!ume i i ~ t i ~ n msd@
!\
j y' /' / i ( i s
Figure 1.163. ModeIs of backarc spreading (Tamaki and Honza, 1991).
(3) The active back-arc extension in the Okinawa Trough and the Taupo Depression in New Zealand can be explained by the injection model. Otofuji (1996) proposed a "double door" opening mode with a fast spreading rate of 21 cm/year for the evolution of the Japan Sea, caused by the injection of asthenosphere with a low viscosity beneath the Japan Sea area. Model 3 is a plate kinematic model. The retreat of a back-arc plate forms a back-arc basin (Dewey, 1980). Model 4 is also a plate kinematic model. The retreat of a fore arc plate forms a back-arc basin. This model seems attractive. Jackson et al. (1975) found the periodicities of rotational motions of the Pacific plate. When the direction of the Pacific plate changed and obliquely subducted, the compressional force of oceanic plate to continental plate decreases. That means that the retreat of fore arc plate occurs.
Chapter 1
230
L NE CHINA UPWELLING
40N
L/ j~jl "~ • ".i I i ,
,
, °
°.
t~,,, e
:.,
•
B I":
140E
20N 120E Figure 1.164. Distribution of Cenozoic basalts and active rift systems in the northeast China region. Arrows indicate the horizontal convective current in the upper mantle associated with the upwelling of the asthenosphere beneath the region. A: Baikal Rift; B: Shanxi Graben; C: Tancheng-Lujiang Fault; D: Okinawa Trough (Tatsumi et aI., 1990).
INTRA-PLATE VOLCANISM CONTINENTAL RIFT SYSTEM P
+
+
÷
•
° E4
+'m,(÷
+'lrff+
l lll {11{
BACKARC BASIN
o,
;
÷
+
VOLCANIC BASIN
. .
°
5rrrvrrrrrrff+
÷
,
.4, y ~ .
~ /~/;~'0"0 Asthen°sPhi re.~q~ Asthenosphere
Upwellingof Asthenosphere ~
f
e,~,~
Figure 1.165. Schematic diagram showing the effect of upwelling of asthenosphere beneath a continent which is adjacent to an arc-trench system (Tatsumi et al., i990).
Miocene-Pliocene Hydrothermal Ore Deposits
231
In any model, back-arc basins form under the extensional stress regime and are associated by Mariana-type subduction by Uyeda and Kanamori (1979) rather than Chilean-type subduction. In this tectonic situation, intense bimodal volcanism and associated seawater circulation occur, resulting to the formation of Kuroko deposits on the seafloor and formation of vein-type mineralization under subaerial condition and intense hydrothermal and volcanic CO2 fluxes to ocean and atmosphere. Such fluxes affect the long-term environmental changes (see Chapter 4).
1.6. Other hydrothermal ore deposits 1.6.1. Polymetallic vein-type deposits The Honshu arc is divided into Northeast and Southwest Honshu by the Tanakura Tectonic Line or the Itoigawa-Shizuoka Line (Fig. 1.166). There are many geological and geochemical differences in these two geologic provinces. Neogene altered volcanic rocks (Green tuff) are widely distributed in Northeast Japan, whereas Cretaceous-Neogene granitic rocks and metamorphic rocks are in Southwest Japan. 87Sr/S6Sr ratio of volcanic rocks and granitic rocks (Shibata and Ishihara, 1979) (Fig. 1.166) and ~348 and sulfur content of Quaternary volcanic rocks (Ueda and Sakai, 1984) are different in two provinces (Fig. 1.167). As for the mineralization at the middle Miocene age in the Northeast Japan, Kuroko and epithermal base-metal veins have been formed. No enrichment of Sn and W is found in these deposits. In contrast, in Southwest Japan, polymetallic veins (so-called xenothermal-type deposits in the sense of Buddington (1935) or subvolcanic hydrothermal type in the sense of Cissartz (1928, 1965) and Schneiderhrhn (1941, 1955) occur. Examples of these deposits are Ashio, Tsugu, Kishu and Obira. All these vein-type deposits have formed at middle Miocene age in western part of Tanakura Tectonic Line under subaerial environment. In these deposits, many base-metal elements (Sn, W, Cu, Pb, Zn) and small amounts of Au and Ag are concentrated. These deposits are associated with felsic volcanic and plutonic rocks along the Median Tectonic Line (MTL) or south of MTL. According to Ishihara (1977), these granitic rocks are ilmenite-series while granitic rocks in Green tuff region are magnetite-series. 3 1.6.1.1. Ashio deposit The Ashio mine which is located in middle Honshu is one of the largest Cu ore producer in Japan until it was shut down in 1973 (Kanehira, 1991). The ore deposits
3Ilmenite-series and magnetite-series granitic rocks are defined as follows (Ishihara, 1977): the magnetiteseries and ilmenite-series granitic rocks are distinguished by the presence or absence, respectively, of magnetite in polished sections.
Chapter 1
232
HOKKAIDO
TANAKURA TECTONIC LINE ,"
0
0
o
o.,d KYUSHU
MEDIAN TECTONIC LINE
d
O U
I
300 KM ,
i
I
Figure 1.166. Distribution or initial 878r/86Sr (r i) for the Cretaceous-Paleogene plutonic rocks. Open squares: gabbros; open circles: granites. Numbers indicate the last two or three digits of ri values. Solid squares and solid circles imply the magnetite-series; open squares and open circles the ilmenite-series. (Shibata and Ishihara, 1979).
formed at middle Miocene age, associated with rhyolitic intrusive activity (Nakamura, 1970). The deposits occur in rhyolitic body as a funnel-shaped mass (Figs. 1.168 and 1.169). The deposits are characterized by conspicuous metal zoning and polymetallic mineralization. From the centre to margin of the mine district, the following zonings are recognized; S n - W - B i - C u zone, Cu-As-Zn zone, and Z n - P b - C u - A s zone (Nakamura, 1970). 3180 values of quartz are +12.8%o to 12.2%o (SMOW) (early-stage ( S n - W - B i Cu) quartz with cassiterite), +11.5%o to +11.4%o (quartz with stannite), +10.3%o to +8.8%0 (middle Cu-As-Zn stage), +9.4%e to +6.4%0 (late Z n - P b - C u - A s stage) (Shimazaki et al, 1993). If temperature is 300°C estimated from homogenization temperatures
Miocene-Pliocene Hydrothermal Ore Deposits 240
WEST JAPAN
420
6
70 60
233
Southwest Japan (SW) • RyukyuArc (RY)
A[
50
v High 834S rock (HI)
RY
* Alkaline (AL) ~" 4O
@
0..
~" 30 20
®
HI
10
+5
501
+10
EAST JAPAN
~ I M
30
20
0
-2
;H
HK
~'
_.~_~
0
+20
o Akita-Komagatake(AK) Hakone (HK) * Nasu zone (NA) , Chokai zone (CH) = Izu-Mariana(IM)
4O
(3.
+1'5
NA
+5
,
+10 ~34s (%,,)
+15
Figure 1.167. Sulfur content v s . 3 3 4 8 value for Quaternary volcanic rocks of Japan. Field bounded by solid lines show two volcanoes (AK and HK), two volcanic zones (NA and CH) in Northeast Japan, three volcanic belts (IM, SW and RY), alkaline rocks (AL) and volcanic rocks of unusually high 334S values (HI) in Ryukyu belt. Symbols surrounded by small circles show 334S values of Satsuma-Iwojima volcanic rocks in West Japan (Ueda and Sakai, 1984).
of fluid inclusion throughout all stages of mineralization, ~180 values of hydrothermal solution are early-stage: +4%° to +6%o; middle stage: +2%o to +3%0; and late-stage: - 1 % o to +2%0 (Shimazaki et al., 1995). This variation indicates that hydrothermal solution at early-stage was dominantly o f magmatic origin (or igneous origin), and a contribution of meteoric water component increased with the stage of mineralization. Main opaque minerals are chalcopyrite, cassiterite, stannite, arsenopyrite, bismuthinite, pyrrhotite and sphalerite. The FeS content of sphalerite is high (about 18 tool% FeS). Dominant gangue minerals are quartz, sericite, kaolinite and siderite. The Fe content of chlorite is very high (Fig. 1.170). All these data suggest low fs2 and f o 2 conditions.
Chapter 1
234 A V
V," V
Jv
V
V V
v
V
V
,
I1[
"","7"
\
/
II1
~°~°~
~_-'f.~" B
I: (Central) Sn-W-Bi-Cu zone I1: (Intermediate) Cu-As-Zn zone Ill: (Marginal) Zn-Pb-Cu-As zone
Figure 1.168. Geologic map of the Ashio mining area, showing the mineral zoning in the Ashio rhyolitic body and the location of the "Kajika" deposits (alter Ashio mine, partly revised) (Nakamura, 1970).
~34S values o f sulfides are close to 0%~ (Shimazaki, 1985), suggesting magmatic or igneous origin and no contribution of seawater sulfur. These values are quite different from those o f Green tuff sulfur (Kuroko and epithermal base-metal vein-type deposits; +2%~ to +7%o).
1.6.1.2. Tsugu deposit The Tsugu g o l d - a n t i m o n y deposit is located in the central part o f Honshu, Japan (Fig. 1.171). Geologically, the Tsugu deposit occurs in the Ryoke metamorphic terrane,
Miocene-Pliocene Hydrothermal Ore Deposits m
1200 A 1000 8O0 600 400
I [ ~~TM,. ~
•
235
\I
~
/
2OO
0t Rhyoliteweldedtuff ['L~
~
Rhyolite
S,a,e
~
OreVei. ' ai,k; depos,,s
lternationof sandstone and slate
Basalbreccia
H : Hotel vein,
lO00m
~Chert
[ q ~ " - I Rhyolite(lithoklite) dike. ~
500
J : Jimbovein, S : Shinseivein, Y : Yokomabuvein,
Figure 1.169. Geologic section along the line A B in Fig. 1.168 showing the mineral zoning in the Ashio rhyolitic body (Nakamura, 1970).
Mg4AI/ /
/"\ \
/ • ', /
/
L/'_ _~1.42 / . . . .
/',,,,,' Mg6
/"\ \
z13\
', ../.
/"k 06 l U ~ Fe(Mn)]4AI2
\,it/
/
"x,~.T12 - IlL/ -
\ 10_4/03\
\~.~/ i \
--$~--2_~
',, .,, ,,, ,,, ,,, ,,, ,,, ,,, \ [Fe(Mn)]6
Figure 1.170. Diagram showing the octahedraI composition of chlorites from the subvolcanic hydrothermal deposits, propylite, and Kuroko deposits in Japan (Nakamura, 1970). Chlorite occurring as a gangue mineral in the subvolcanic hydrothermal deposits: Nos. 1, 2, 3 and 4: Chlorite from the Ashio copper mine. Nos. 5, 6, and 7: Chlorite from the Kishu mine. No. 8: Chlorite from the Arakawa mine. Nos. 9 and 10: Chlorite from the Ani mine. No. 11: Chlorite from the Osarizawa mine. Chlorite from the so-called propylite: No. 12: Chlorite from the Yugashima mine. No. 13: Chlorite from the Budo mine. Chlorite from the Kuroko deposits: No. 14: Chlorite from the Wanibuchi mine.
adjacent to the outer z o n e o f southwestern Japan. In this z o n e and in the R y o k e m e t a m o r phic terrane, A u deposits are rare, w h e r e a s there are m a n y H g - S b deposits. O n the other hand, as already noted, in the Tertiary s u b m a r i n e v o l c a n i c r e g i o n ( G r e e n tuff region) and Quaternary v o l c a n i c region, m a n y e p i t h e r m a l v e i n - t y p e deposits occur. M i n e s w h i c h h a v e p r o d u c e d b o t h Sb and A u are very rare. T h e r e f o r e , the T s u g u deposit is an unusual type o f A u deposit in Japan. I g n e o u s activity related to the m i n e r a l i z a t i o n indicates that the age o f m i n e r a l i z a t i o n was m i d d l e M i o c e n e and this age is different f r o m that o f e p i t h e r m a l A u - A g mineralization, m o s t l y late M i o c e n e to Pleistocene, as already noted.
Chapter 1
236
17= z,0o - - -
35 °
~.
Tsugu D e p o s i t
i 133°
0
200kin
138°
Figure 1.171. Map showing location of the Tsugu gold-antimony deposit. I: Green-tuff region, 2: outer zone of southwestern Japan, TTL: Tanakura tectonic line, 1STL: Itoigawa-Shizuoka tectonic line, MTL: Median tectonic line (Shikazono and Shimizu, 1988a).
The previous studies on the Tsugu deposit (Tsuboya, 1936; Tatsumi, 1948; Shikazono and Shimizu, 1988b) demonstrated that: (1) Opaque minerals include stibnite, jamesonite, cinnabar, gold, pyrite, pyrrhotite, arsenopyrite, marcasite, sphalerite, galena and chalcopyrite. (2) The ore minerals display a zonal distribution: gold and cinnabar are enriched in the upper part of the veins, and sphalerite, galena and chalcopyrite are more abundant in the deeper parts. Pyrrhotite and arsenopyrite are distributed throughout the veins. (3) From the mode of occurrence of opaque minerals it is considered that pyrrhotite and sphalerite were precipitated at an early-stage, gold, pyrite, marcasite, stibnite and cinnabar were precipitated at a late-stage, and arsenopyrite was precipitated throughout the mineralization period. (4) The accessory minerals are tourmaline and apatite. Hg content of electrum ranges from nil to 8.5 wt%, and Ag content ranges from
237
Miocene-Pliocene Hydrothermal Ore Deposits Au
(z-solid / so, t,o
j
•
•
•
Ag
Hg 10
20
30
40
50
60
70
80
90
Hg atomic % Figure 1.172. Chemical composition of mercurian gold plotted on a Au-Ag-Hg triangular diagram. Data sources: Tsugu (open square; Shikazono and Shimizu, 1988b); L~ngsele (solid circle; Nysten, 1986); Aitik, northern Sweden (solid triangle; Nysten, 1986); Kazakhstan (cross; Naz'mova and Spiridonov, 1979); Bajupura-Dariba (solid square; Basu et al., 1981)• Phase boundaries (solid lines) at 450°C were drawn from Basu et al. (1981) (Shikazono and Shimizu, 1988b).
6.0 to t 1.5 wt%. The chemical compositions of gold from the Tsugu deposit are plotted in terms of Au, Ag and Hg, together with data from the literature (Fig. 1.172). The Au/Ag ratios of electrum from the Tsugu deposit are high compared with those from the other mercurian gold occurrences. Note that Hg content varies inversely with Au content (Fig. 1.173), implying that Hg substitutes for Au. A frequency histogram of the Ag content of electrum from epithermal gold-silver vein-type deposits and the Tsugu deposit (Fig. 1.174) clearly indicates that the Au/Ag of electrum from the Tsugu deposit is higher than that from epithermal vein-type deposits. Electrum and cinnabar occur in the same part of the vein, suggesting that these minerals were in equilibrium. The following reaction can be used to determine which factors are important for controlling mercury content of electrum in equilibrium with cinnabar: (Hg) + 1/2S2 = HgS
(1-78)
where (Hg) denotes the Hg component in electrum. From the equilibrium relation for reaction (1-78), we obtain, NHg
1/(gl-78fs2 YHg) 1/2
(1-79)
where NHg is the atomic fraction of the Hg component in electrum, K1-78 is the equilibrium constant of reaction (1-78), i s 2 is the sulfur fugacity, and gHg is the activity coefficient of the Hg component in electrum.
238
Chapter 1 8
6
E
4
-r
2
0
70
75
80
85
90
Au atomic % Figure 1.173. Relationship between mercury and gold contents of mercurian gold from the Tsugu goldantimony deposit (Shikazono and Shimizu, 1988b).
From (1-79) it is clear that the Hg content of electrum is related to f s 2 and K1-78. Using the thermochemical data for reaction (1-78) by Barton and Skinner (1979), isoactivity lines for Hg in electrum may be drawn on a log fsz-temperature diagram (Fig. 1.175). At a given temperature, the activity of Hg in electrum increases with a decrease in fs2. Therefore, the f s 2 for the mercurian gold of the Tsugu deposit is inferred to be relatively low. Microprobe analyses of sphalerite indicate that sphalerite from the Tsugu deposit is markedly higher in iron than that from the epithermal Au-Ag vein-type deposits in Japan. Measurements of homogenization temperatures of fluid inclusions that were obtained from quartz coexisting with sphalerite, pyrrhotite and chalcopyrite of the earlystage of sulfide mineralization, and from quartz coexisting with electrum, arsenopyrite, pyrite and sericite of the late-stage of mineralization. Homogenization temperatures of fluid inclusions are 281-345°C for the early sulfide stage, and 275-295°C for the latestage of gold mineralization. These fluid inclusion data, stability fields of arsenopyrite, pyrite and pyrrhotite, and the iron content of sphalerite can define the possible ranges of sulfur fugacity and temperature for the Tsugu mineralization (Fig. I. 176). The ranges of is2 and temperature for epithermal Au-Ag vein-type deposits in Japan have been clearly defined based on the chemical composition of sphalerite and electrum, and homogenization temperatures of fluid inclusions (Shikazono, 1985d). Values of fs2 for the Tsugu deposit are lower than the typical ranges of values for the epithermal Au-Ag vein-type deposits in Japan (Fig. 1.176). Such a low i s 2 is in accord with the high Hg content of electrum in the Tsugu deposit.
Miocene-Pliocene Hydrothermal Ore Deposits
239
70 ¸
60
50 O ©
40
(1) ~4 30
20
10
20
40
60
Ag a t o m i c
80
I00
%
Figure 1.174. Frequency (number of analyses) histogram for Ag (atomic %) of gold from the Tsugu deposit (solid) and epithermal gold-silver vein-type deposits in Japan (open). I: sampte I; II: sample II. Data sources: Tsugu deposit (Shikazono and Shimizu, 1988b); epithermal gold-silver vein-type deposits in Japan (Shikazono 198l, 1986; Shikazono and Shimizu, 1988b).
~xx%
o .a -15
150
200
250
300 Temp.(°C)
Figure 1.175. Activity of S2-temperature diagram showing iso-Hg contents contours for gold. The calculations were carried out using thermochemical data of Craig and Barton (1973).
Chapter 1
240
/-
~
-10
'
J
'
o -
/
c
//////" ,
,
,
r
150
,
,
,
,
I
200
~
I
250
I
i
t
i
I
300
.
e
%t-.
,. ~,.o-
//...~,%.-'//
O k/~
J
i
-1 ~
i
-
i
i
i
i
1
350 Temp,(°C)
Figure 1.176. Possible ranges of activity of $2 (as2) and temperature for the Tsugu deposit (solid and dotted areas) and epithermal gold-silver vein-type deposits in Japan (shaded area) (Shikazono and Shimizu, 1988a).
Mercurian gold has been reported from several Au deposits: Longsele, Aitik, Kazakhstan, and Bajpura-Dariba (Naz'mova and Spiridonov, 19797 Basu et al., 1981, Nysten, 1986). The nature of the opaque minerals coexisting with mercurial gold in these minerals are pyrrhotite, galena, sphalerite, stibnite, and jamesonite, and less common are cubanite, arsenopyrite, pyrite, and owyheeite. Pyrrhotite, cubanite, and arsenopyrite suggest relatively low fs2, but the temperatures of formation of these deposits have not been estimated. 1.6.1.3. Kishu deposit The Kishu deposit, located at middle Honshu, is associated with Kumano acidic rocks intruded into the Shimanto and Kumano formations. K-Ar age determination indicates that the activity of Kumano acidic rocks occurred at middle Miocene (144-1 m.y.) (Ikebe, 1973). The stage of ore mineralization from early to late is (1) Cu-Fe (chalcopyrite, pyrite, chlorite, quartz), (2) Cu-Zn-Pb (chalocopyrite, sphalerite, galena, sericite, quartz, siderite, calcite, fluorite, chlorite), (3) Au-Ag (native gold, argentite, pyrite, sphalerite, chalcopyrite, chlorite, quartz, adularia, sericite), (4) calcite (Nakamura and Miyahisa, 1976). ~348 of sulfides from Kishu is very low (-14.2%o) (Shimazaki, 1985), indicating sedimentary sulfide sulfur source. 1.6.1.4. Obira deposit Obira deposit is located at south Kyushu. It occurs in Shimanto Group rocks composed of shale, and limestone, associated with middle Miocene granitic intrusive rocks. Polyascendant zoning and mineralization are observed in the Obira deposit. From early- to late-stage mineralizations are:
Miocene-Pliocene Hydrothermal Ore Deposits
241
cassiterite-apatite-tourmaline --+ cassiterite-wolframite-molybdenite-pyrite --+ pyrite-arsen•pyrite-pyrrh•tite-magnetite-cha•c•pyrite-stannite-bismuthinite-ga•enatetrahedrite --+ marcasite-pyrite --+ marcasite-Sb minerals (stibnite, jamesonite, berthierite, boulangerite). Occurrence of pyrrhotite, arsenopyrite and high iron content of sphalerite indicate low i s 2 and / O 2 conditions. Decrepitation temperatures of fluid inclusions in quartz from S n - C u - A s stage are 403-322°C (Enjoji and Takenouchi, 1976). ~34S values of sulfides are low (-6.7%o) (Shimazaki, 1985), suggesting a contribution of sedimentary sulfide sulfur. Distinct features of the above-mentioned polymetallic deposits are summarized as follows: (1) The FeS content of sphalerite and Fe content of chlorite from these deposits are high. (2) Temperature of formations is high and variable. (3) fs2 and f02 are low. (4) Many metals (Cu, Pb, Zn, Au, Ag, Sn, W, Bi) have been produced. (5) Orebody and regional metal zonings are conspicuous. (6) Salinity of inclusion fluids varies widely from high to low. (7) Common gangue minerals are kaolinite and sericite, but K-feldspar is not found, suggesting low pH of ore fluids. (8) The deposits have been formed under the subaerial environment. No contribution of seawater and low fs2 and f02 conditions are consistent with the geologic environment.
1.6.1.5. Temperature and sulfur fugacity estimated from iron and zinc partitioning between coexisting stannite and sphalerite and coexisting stannoidite and sphalerite Stannite is the most common tin sulfide mineral in the ore deposits associated with tin mineralization. This mineral sometimes contains appreciable amounts of zinc, together with iron. Several workers have suggested that the zinc and iron contents of stannite are related to temperature. With respect to the study of the phase relationships in the pseudobinary stannite-kesterite system, Springer (1972) proposed zincic stannite as a possible geothermometer mainly based on the chemical compositions of the two exsolved phases (stannite and kesterite). Nekrasov et al. (1979) and Nakamura and Shima (1982) experimentally determined the temperature dependency of iron and zinc partitioning between stannite and sphalerite. Because natural stannite contains a considerable amount of zinc, sphalerite contains a considerable amount of iron, and these contents can be easily analyzed using an electron microprobe, a stannite-sphalerite pair is expected to be a useful indicator of formation temperature and sulfur fugacity. Iron and zinc partitioning between stannite and sphalerite is represented by the exchange reaction. Cu2FeSnS4 (in stannite)+ ZnS (in sphalerite) = Cu2ZnSnS4 (in stannite)+ FeS (in sphalerite)
(1-80)
Chapter 1
242
"" 1.0 cn o
I .
~.
1
Nekrasov el 0t.0979),
looKd =l.27&.lOd'T-1-l.174
7 Nokomuro & Shimo ( 1982 ), log Kcl = ZS.103.T- 1_3fi
250
2.0
300
350
l~a
400 450
1.6
590~C
~4
f.2 103"T -1
Figure 1.177. Comparison between the stannite-sphalerite geothermometer after Nekrasov et al. (1979) and one after Nakamuraand Shima (1982). Crossbars indicate experimental uncertainties (Shimizu and Shikazono, 1985). 153
The partition coefficient, Kd, for the above reaction is expressed by,
Kd = (Xcu2FeSns4/Xcu2znsnS4)stannite/(XFeS/XZnS)sphalerite
(1-81)
where X denotes mole fraction of a given component in stannite and sphalerite. Nekrasov et al. (1979) and Nakamura and Shima (1982) reported a temperature dependency of iron and zinc partitioning between stannite and sphalerite (Fig. 1.177) logKd = 1.274 x 103T -I - 1.174 (Nekrasov et al., 1979)
(1-82)
logKd = 2.8103T -1 - 3.5 (Nakamura and Shima, 1982)
(1-83)
Both geothermometers are in agreement being close 380°C (Fig. 1.177). However, at the lower and higher temperatures the difference between the temperatures estimated from equations (1-82) and (1-83) becomes larger. Figure 1.178 represents a comparison between the stannite-sphalerite temperatures and homogenization temperatures of fluid inclusions or sulfur isotope temperatures. It can be seen in Fig. 1.178 that Nakamura and Shima's geothermometer would be rather consistent with the temperature estimated based on the fluid inclusions or sulfur isotope studies. It is notable that almost all stannite-sphalerite temperatures are within 30°C of average homogenization temperatures and sulfur isotope temperatures. Temperatures estimated based on equations (1-82 and 1-83) range from ca. 250°C to 350°C for both skarn-type and vein-type deposits except the Yatani epithermal Au-Ag vein-type deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
243
°C
°C
400.
400 /
E O
IoT ~K
or'
E 300,
(N
tO
200
200
100
100
@
C O D
"6 2
Nekrasov et al (1979) 0
Nakamura and Shima (1982) t
0 0
100
200
300 °C
100
200
i
300 °C
stannite-sphalerite geothermomentry Figure 1.178. Comparison between the stannite-sphalerite temperatures and fiIIing temperatures of fluid inclusions or sulfur isotope temperatures.NT: Nakatatsu, OB: Obira, KN: Kant, KG: Kuga, TM: Tsumo, KM: Kamioka, OT: Ohtani, KU: Kaneuchi, Ak: Akenobe, TT: Takatori,YT: Yatani (Shimizu and Shikazono, 1985).
The chemical compositions of the coexisting stannite and sphalerite are plotted on log(XFeS/XZnS)sphaterite--log(Xcu2Fe2SnSg/XCu2ZnSnS4)stannite diagrams (Figs. 1.179 and 1.t80). Iso-logfs2 lines on these diagrams are based on equation (1-83), and thermochemical data on the F e - Z n - S system by Scott and Barnes (1971). The line 20.8 mol% FeS in sphalerite, which corresponds to the composition of sphalerite in equilibrium with pyrite and pyrrhotite at 1 bar as determined by Boorman (1967), is also given on the diagrams. It is evident in Figs. 1.179 and 1.180 that the FeS content of sphalerite coexisting with pyrite is generally lower than that of sphalerite coexisting with pyrrhotite or both pyrite and pyrrhotite. Iso-FeS content lines for sphalerite in equilibrium with pyrite or pyrrhotite were drawn on the log fs2-temperature diagram (Figs. 1.179 and 1.180) using thermochemical data by Scott and Barnes (1971) and Barton and Skinner (1979). The relationship between the iron content of stannite in equilibrium with sphalerite and pyrite or with sphalerite and pyrrhotite was derived based on thermochemical data by Scott and Barnes (1971), Barton and Skinner (1979) and Nakamura and Shima (1982). These types of deposits are skarn-type polymetallic (Sn, W, Cu, Zn, Pb, Au, Ag) vein-type and S n - W vein-type deposits. As shown in Fig. 1.181, the fs2-temperature range for each type of deposits is different; at a given temperature, fs2 increases from S n - W vein-type through skarn-type to polymetallic vein-type deposits. It is interesting to note
Chapter 1
244
250°C
¥r
].
Ti.0
am
/ ~ 1 2 C U')
u7 rt/~
U
(D
450°C
0
0
-t
-t
0
log XFeS in spholerite XZnS Figure 1.179. log(XFeS/XZnS)sphalerite--Iog(XCuzFeSnS4/Xcu2ZnSnN4)stannite diagram showing that sphalerite and stannite are associated with pyrrhotite (Po) and/or pyrite (Py). Temperature lines are based on data by Nakamura and Shima (I982). Solid curves show logfsz based on data by Scott and Barnes (1971) in the pyrrhotite field. Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).
that i s 2 associated with gold-silver mineralization in polymetallic deposits is relatively high compared with those of the other types of deposits. This difference is similar to that found in epithermal vein-type deposits; fs2 of epithermal A u - A g veins is higher than that of base metal veins. Shimizu and Shikazono (1987) studied the compositional relations of coexisting stannoidite, sphalerite and tennantite-tetrahedrite (Fig. 1.182). Based on these data they estimated the sulfur fugacity of stannoidite-bearing tin ore. Considering the complementary work on stannite-bearing tin ores from Japanese ore deposits (Shimizu and Shikazono, 1985), a comparison between environmental conditions of these two types of tin sulfides was made. Their study is described below. The chemical compositions of coexisting sphalerite and tennantite-tetrahedrite from the mines were determined. Except the Ashio polymetallic deposits, the other deposits have been formed at late Cretaceous related to felsic magmatism. Scanning patterns and many analytical data obtained by electron-microprobe analyzer reveal that most stannoidite grains are compositionally homogeneous. There is
Miocene-Pliocene Hydrothermal Ore Deposits
245
YT
250oc 9'/', ,
/iK@iL,
I C 0
/
3sooc
,..5ooc
C
o
,
/
-1 -2
.
Py
1
tog XFeS
in sphalerite
XZnS Figure 1.180. log(XFeS/XZnS)sphalerite Iog(Xcu2FeSnS4/XCu2ZnSnS4)stannitediagram showing each deposit is skarn-type or vein-type. Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).
a wide range in extent of Fe and Zn substitution in stannoidite (NFe/l~Zn between 2.03 and 14.4). The stannoidite from the Tada deposit is the richest in Zn (5.24 wt%, Zn, 9.09 wt% Fe) and has the approximate formula Cu8Fe2ZnSn2S12; that from the Konjo deposit is the richest in Fe (11.90 wt% Fe, 1.69 wt% Zn) and has the formula Cu8Fe2(Fe0.77Zn0.19)Sn2S12. From the Fe2+/Zn 2+ ratio of stannoidite, a continuous solid solution is inferred to exist between Cu8Fe2ZnSn2S12 and CusFe2FeSn2SI2. Iron and manganese contents of sphalerite coexisting with stannoidite are in the range from 0.17 to 1.93 wt%, and from 0.02 to 0.16 wt%, respectively. Lee et al. (1974) conducted an experimental study on the equilibrium for the assemblage of stannoidite-chalcopyrite-bornite-mawsonite-S2 (gas) in a temperature range from 430 to 300°C. Curves A and B in Fig. 1.183 correspond to fsz-temperature relationships for this equilibrium assemblage for aFe = 1 and aFe = 0.1, where aFe is the activity of the Cu8Fe2FeSn2S12 component in stannoidite solid solution. As mentioned already, Shimizu and Shikazono (1985) have estimated the is 2temperature range for stannite-bearing assemblages from Japanese vein-type and skarntype tin deposits. This estimated fs2-temperature region is also shown in Fig. 1.183. The fsz-temperature range for the formation of these two types of tin sulfides is different.
246
Chapter 1
Cn 0 I
10
Temperoture, *C 11
250 30O 35O *C Figure 1.181. Temperature-log fs: diagram. Skarn-type deposits (solid squares) are considered to be formed under lower fs2 condition than vein-type deposits (open squares). Abbreviations are the same as in Fig. 1.178 (Shimizu and Shikazono, 1985).
Microscopic observation suggests that stannite formed earlier than stannoidite on the scale of one polished section, although in general coexistence of stannoidite and stannite cannot be observed in the same polished section. If stannite formed at an earlier stage than stannoidite, it could be inferred that fs2 increased or that temperature decreased (or both) with evolution of tin mineralization. The Fe2+/Zn 2+ ratio of coexisting stannoidite, sphalerite and tennantite-tetrahedrite from the Tada, Omodani and Ohmidani deposits is low, compared with that from the other deposits such as the Ashio, Akenobe and Ikuno deposits (Fig. 1.182). The Tada, Omodani and Ohmidani deposits are characterized by polymetallic (Zn-Cu-Ag-Au; Zndominated) mineralization, and tungsten is not recovered from these deposits. On the other hand, the Akenobe, Ikuno and Ashio deposits are characterized by polymetallic (Cu-AnPb-Sn-W-Ag-Au-Bi) mineralization, and tungsten is recovered from these deposits. The Fe/Zn ratio of coexisting stannoidite, sphalerite and tennantite from the tungsten-bearing polymetallic deposits is higher than in those from the tungsten-free polymetallic deposits. High content of indium in chalcopyrite from tungsten-beating polymetallic deposits probably indicates a high temperature of formation. Sphalerite rarely appears where the Fe2+/Zn 2+ ratio of the coexisting stannoidite exceeds approximately 1. These differences imply that the tungsten-bearing deposits formed under lower fs2 or higher temperature conditions (or both) than tungsten-free deposits.
Miocene-Pliocene Hydrothermal Ore Deposits
247
KONJO
2-
2
I~KUNO A ¢N
ASHIO
1
~OMO_~I
ASHIO N'ENOE
AKENOBEI
~,
/ FUKOKU
=~"KOKU
tOMO' ~DANI OHMIDANI 0
I 2 3 4 ( Fe 2" I Zn 2" )sp x 10 2
0
1 2 3 ( F e z. / Zn 2" )tenn-tetr
Fig 1.182. Fe2+/Zn2+ of sphalerite (left) and of tennantite-tetrahedrite-series mineral (right) as a function of FeZ+/Zn2+ in stannoidite (atomic proportions) (Shimizu and Shikazono, 1987).
Figure 1.183 shows i s 2 and temperature region for stannoidite-bearing tin deposits is different from that for stannite-bearing tin deposits.
1.6.2. Hg and Sb deposits Hg deposits are distributed along the MTL (Median-Tectonic Line) in Southwest Japan and in Northeast Hokkaido (Fig. 1.184). The deposits are vein or disseminated in form. The deposits are hosted by sedimentary and igneous rocks. No K - A r age data on the deposits are available. However, from the age of host rocks, the age of igneous activities along MTL and the studies on the movement of MTL and K - A r ages of Au-veins associated with Hg mineralization in Northeast Hokkaido (e.g., Khonomai), it is likely that these deposits formed at middle Miocene age. However, mercury mineralization in Kitami Province (north Hokkaido) occurred at approximately the same age as the epithermal gold-silver mineralization in the same district (4.5-5.3 Ma) (Maeda, 1997). Main ore minerals are few. They are cinnabar, metacinnabar, stibnite and pyrite. Gangue minerals are quartz and calcite. Fluid inclusion decrepitation temperatures
248
Chapter 1
0001
Stcmnoidite r e g i o n / 0005
~'-. . .
0
jf
10 ," /// B
Stotnnite r e g i o n (Shimizu & 5hikazono 1985)
./
15
A
"
Temperature 20
200
250
3 0
350"C
Figure 1.i83. Temperature-log fs2 diagram for stannoidite-bearing ores from Japanese veto-typedeposits. MW = Mawsonite, Sd = Stannoidite, Bn = Bornite, Cp = Chalcopyrite(Shimizu and Shikazono, 1987).
for quartz and calcite from the Itomuka Hg deposit are 328-337°C and 201-242°C, respectively (Enjoji and Takenouchi, 1976). ~34S of cinnabar and stibnite are wide in a range from -16%e to +15%o, and from -15%o to +4%0 for cinnabar and metacinnabar, and stibnite, respectively (Shikazono and Shimizu, unpublished). Usually, 334S values of cinnabar and stibnite are low less than 0%o except for these minerals from Green tuff region. These low 334S values imply that sulfide sulfur of Hg and Sb deposits came dominantly from sedimentary sulfur, and no contribution of seawater sulfur. The sulfur isotopic data are consistent with geologic environments of Hg and Sb deposits; Sedimentary rocks are dominant and marine rocks are not present in Sb-Hg mineralization districts. However, a few samples of stibnite and cinnabar from the deposits in Green tuff region display high ~34S values. In contrast of this interpretation on the origin of sulfur, Ishihara and Sasaki (1994) thought that sulfur came from ilmenite-series granitic rocks. However, these rocks are not found in the north Hokkaido. Horikoshi (1995) showed areal extension of volcanic belts since middle Miocene in the Japanese Islands. He suggested that Hg and Sb mineralization in outer zone of Southwest Japan and north Hokkaido (Kitami district) related to forearc igneous activities in trench side. ~34S values become heavier across the Northeast Japan arc from the trench to the back-arc side. He thought that the changes in tectonic character from forearc to arc environment controlled the ~34S values.
Miocene-Pliocene Hydrothermal Ore Deposits
249
p R
i Sado
O0
,
Sanin Kyoto s
g
% o okue~_ ~
~o~s p~o.4\~ce
e~ 're'=
\\ra
zone,
0 I
I
I
3 0 0 km I
qxl'l ,~*~"
o~te~
,0
• e • e o
Hg Ore Deposits Sb Ore Deposits Au-Ag Ore Deposits Pb-Zn Ore Deposits Sn-W Ore Deposits
o Vein o
m Skarn
D
Figure 1.184. Distribution of major sulfide deposits in Miocene volcanic and plutonic provinces. Class i and 2 are ore deposits (that is, the largest and second largest are shown together with the larger symbols). Class 3 is shown as a smaller circle and box. Tin ores occur associated with base-metal deposits at Mt. Okue, and tungsten deposits are concentrated in Yakushima, Kyushu (Ishihara and Sasaki, 1991).
1.6.3. Gold-quartz vein-type deposits (mesothermal-hypothermal vein-type deposits) 1.6.3.1. Geology, mineralogy and geochemistry G o l d - q u a r t z v e i n - t y p e ( m e s o t h e r m a l - t y p e a n d h y p o t h e r m a l - t y p e d e p o s i t s in t h e s e n s e o f L i n d g r e n ( 1 9 2 8 ) ) o c c u r in s e d i m e n t a r y t e r r a n e a s s o c i a t e d w i t h C r e t a c e o u s f e l s i c
Chapter 1
250
0
100
200km i
40° Nishizawa.~°h °
k
/ Hayochine ~Z~;i!~v//Owashi t - - - % ~ ~//Nojiri u ~ Shishiori
K i n k e i \ F::~::/\ ~ ~ i ' )
x AikGwa hio owo
o Amo---------~E~r,', ~ ~ _ 35°
~3 0
_/ .....
133°
/__..~
~~b'~--.
:~; \ ~
~
Suwa
AShiyasu
138°
Figure 1.185. Locations of hypo/mesothermal vein-type deposits in Japan. Abbreviations are the same as those in Fig. 1.62. Open circle: hypo/mesothermal Au vein-type deposits. Solid circle: hypo/mesothermal po]ymetallic vein-type deposits (Shikazono and Shimizu, 1988a).
plutonic rocks or in regionally metamorphosed rocks (Fig. 1.185). Deposits occur along the Tectonic lines such as Hidaka (Hokkaido), Kitakami (northeast Honshu), Tanakura (middle Honshu) and Itoigawa-Shizuoka (middle Honshu) Lines (Fig. 1.185). The deposits occurring in the sedimentary rocks (mainly black shale) are distributed mainly in three districts: Kitakami, Yamizo and Koma (Fig. 1.185). The few deposits in the metamorphic region are the Suwa, Kinkei, Amo and Hashidate (Fig. 1.185). Ages of mineralization in the Hidaka and Kitakami regions may be Cretaceous, considering the ages of associated granitic rocks. Isozaki and Maruyama (1990) thought that the Tanakura Tectonic Line moved at middle Miocene (ca. 15 Ma) related to the opening of Japan Sea. The time of movement of Itoigawa-Shizuoka Line is thought to be also middle Miocene. Ages of granitic rocks near the itoigawa-Shizuoka Line are middle Miocene (Shimizu, personal communication, 1998). These lines of evidence suggest that ages of mineralization along the Tanakura and Itoigawa Tectonic Lines are probably middle Miocene.
Miocene-Pliocene Hydrothermal Ore Deposits
251
AURIFEROUS VEIN Type 1
F7~J
60 Type 2
50
~
f°
after Nedachi (1974) after Yarnaoka (1981)
GOLD-SILVER VEIN E~ after Shikazono (1981) >, 40 O c-
$ u_ 30
20
10
0
0
20
40
60
80
100
Ag atomic % Figure 1.186. Frequency (number of analyses) histogram of Ag content (atomic %) of auriferous vein and gold-silver vein deposits in Japan. Frequency means numbers of analyses (Shikazono and Shimizu, 1987).
Main opaque minerals include native gold, electrum, pyrite, pyrrhotite, chalcopyrite, cubanite, sphalerite, arsenopyrite and tellurobismutite. The amounts of these sulfide minerals are poor, compared with those in epithermal A u - A g vein-type deposits. It is noteworthy that silver minerals are abundant in epithermal A u - A g vein-type deposits, whereas they are poor in gold-quartz veins. The Ag content of electrum is very low (Fig. 1.186) and FeS content of sphalerite is high (6-17 FeS mol%) (Fig. 1.187) (Shikazono and Shimizu, 1987). Combining these compositional data with homogenization temperatures of fluid inclusions, fs2 of ore fluids was estimated (Fig. 1.188). Estimated fs2 range is lower than that of epithermal A u - A g vein ore fluids. Shikazono and Shimizu (1987) calculated Z A u / I ; A g of ore fluids using fluid inclusion data (salinity, temperature), pH, mH2s and mK+. The procedure for estimating Z A u / Z A g by them is given below.
Chapter 1
252
100
•
AURIFEROUSVEIN
[--] GOLD-SILVERVEIN 80
O ¢-
60 t:r
1.1_ 401
20
--
0
2
4
6
8
10 12 14 16 18
FeS mole % of sphalerite Figure 1.187. Frequency (number of analyses) histogram of FeS content (mol %) of sphalerite from auriferous vein deposits in Japan (Shikazono and Shimizu, 1987).
^~,'-~*~//7..: . ~
~,
-10 d
-15 ,
~O~:5,,,,,.,~ ~
150
~' S,4~.,~
~'3P''f:££'O'a
200
250
300
350
Temp (°C)
Figure 1.I88. Typical sulfur activity and temperature ranges for Japanese auriferous vein (dotted) and goldsilver vein (hatched) deposits. Iso-FeS content curves for sphalerite were drawn based on the equation of Barton and Skinner (1979). py: pyrite, po: pyrrhotite (Shikazono and Shimizu, 1987).
It is essential to know the mode of transport o f Au and Ag in ore fluids to consider the factors which control the A g / A u ratio o f native gold and electrum. Many studies on Au and Ag complexes in ore fluids have been conducted and reviewed by several workers (Barnes and Czamanske, 1967; Barnes, 1979; Seward, 1981; Shenberger, 1986).
Miocene-Pliocene Hydrothermal Ore Deposits
253
According to these previous studies, the most dominant dissolved states of Au and Ag in ore fluids are considered to be bisulfide and chloride complexes, depending on the chemistry of ore fluid (salinity, pH, redox state, etc.). However, very few experimental studies of Au solubility due to chloride complex and Ag solubility due to bisulfide complexes under hydrothermal conditions of interest here have been conducted. Thus, it is difficult to evaluate the effects of these important species on the Ag/Au of native gold and electrum. Other Au and Ag complexes with tellurium, selenium, bismuth, antimony, and arsenic may be stable in ore fluids but are not taken into account here due to the lack of thermochemical data. Assuming that Au is transported dominantly as bisulfide or chloride complexes, the following reaction can be used to determine which species are dominant under near neutral conditions• AuC12 + 2H2S -----Au(HS) 2 + 2C1- + 2H +
(1-84)
The equilibrium relation for reaction (1-84) is expressed as 2 2 2 aauci 2/aau(HS) ~ = mauci 2/mAu(HS) ~ = (aci- all+)/(K1-84aH2S)
(1-85)
where a is activity, m is molality, and K1-84 is the equilibrium constant for equation (1-84)• It is assumed that the activity coefficient ratio, YauC12/yAu(US)2, is one. Several investigations to elucidate the geochemical environment of Japanese epithermal A u - A g vein-type deposits have bee,, ,'onducted (Shikazono, 1974a, 1977a, 1978b, 1985a,b). Based on these studies, the valu306of mauci-/mAu(HS) as a function of tem• perature may be calculated (Fig. 1.189), by taking the average values for a c v , all+, au2s, and temperature. The ratio of mAuci- /mAu(HS) increases with increasing temperature and acl (Fig • 1" 189), while an increase m pH causes a decrease in m I-A u t~. / m .~ U. [.r l.~.) ~. ratio " ~l 2 It is apparent in Fig. 1.190 that Au bisulfide species are more abundant than Au chloride species under the conditions common for ore fluids responsible for Japanese A u - A g veins. However, Au chloride species may dominate Au bisulfide species in ore fluids responsible for the gold-quartz (auriferous) vein deposits, as shown in Fig. 1.189. Assuming that AgC12 and Au(HS)~- are the predominant Ag and Au species, the following reaction may be used to derive the relationship between the Ag/Au ratio of native gold or electrum and temperature or other variables: •
•
.
.
2
.
2
2
(Au) + AgC12 + 2H2S = (Ag) + Au(HS)~ + 2C1- + 2H +
2
.
•
(1-86)
where (Au) and (Ag) are the Au and Ag components of native gold and electrum, respectively. From the equilibrium relation for equation (1-86) we obtain,
aAg/aAu ----(a22smagc12 KI-86)/mau(Hs)2a21 a2+
(1-87)
where K1-86 is the equilibrium constant for equation (1-86)• Equation (1-87) implies that the aAg/aAu of native gold and electrum is controlled by temperature, an.-, aN-S, pH, and m -A g t A~.2 / m A. .u t.t a.~.) 2. " ~1 2 The curves representing the relationship between temperature and aAg/aAu (aag, activity of Ag component in native gold or electrum; aAu, activity of Au component in native gold or electrum) are shown in Fig. 1.190. It is assumed that aK+/aN+ is controlled
Chapter 1
254
6 5
t~
3
iI::::: : ' ' ~ : : : :"::::::::' :": : :
z
£
fi
T-,
0
......
131 O --
-2
-3 -4 -5 ,
200
,
,
.
,
250
.
.
.
.
i
300
T(oC )
Figure 1.I89. The relationship between m .... /m . . . . . . . and temperature• Hatched and dotted areas represent • • , AUk.12 . AU[IJ~)2 the probable geochemical envlronment for typical Japanese gold-silver vein and auriferous vein deposits, respectively• A, mcl = 10, mK+ = 2, aH2S = l0 3, K-feldspar/K-mica/quartz equilibrium; B, reel = 1, mK+ = 0.2, aH2S = 10 -35, K-feldspar/K-mica/quartz equilibrium; C, mci = I, inK+ = 0•2, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium; D, mcl- = 0.2, inK+ = 0•04, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium; E, mcl = 0.2, inK+ = 0.04, aH2S = 10 -3, K-feldspar/K-mica/quartz equilibrium; F, mcl = 0.2, mK+ = 0.04, aH2S = 10 -2, K-feldspar/K-mica/quartz equilibrium• Thermochemical data for the calculations were taken from Helgeson (1969), Seward (1973), Drummond (1981), and Henley et al. (I984)• (Shikazono and Shimizu, 1987).
by the K-feldspar-K-mica-quartz assemblage which commonly occurs in Au-Ag vein deposits in Japan (Shikazono, 1974a). The values of HzS activity in ore fluids responsible for A u - A g veins are assumed to be 10 - 2 - 1 0 - 3 , which are estimated from sulfide mineral assemblage and chemical compositions of minerals (e.g., Fe content of sphalerite; e.g., Shikazono, 1974a). The transport of Au in the ore fluids responsible for the gold-quartz veins probably takes place as Au chloride complexes rather than as Au bisulfide complexes (Fig. 1.190). This suggestion is based on the fact that the temperature of formation of gold-quartz vein deposits is high (probably 250-350°C) and the chloride concentration of the ore fluids is high (probably more than 1 molal and less than 10 molal) based on the fluid inclusion studies (e.g., Nedachi, 1974; Shikazono, unpublished). Based on the experimental data on the solubility of Au due to chloride complexes (Henley, 1984) Au is highly soluble at high temperatures and/or in fluids of high chloride concentration. In fluids where precious metal transport is dominated by AgCI 2 and AuC12, the following reaction may be written to consider the compositional variation of native gold or electrum: (Au) + AgCI~- = (Ag) -t- AuCI~-
(1-88)
Miocene-Pliocene Hydrothermal Ore Deposits 5
255
A
4 3
C
~2
%1 t~
~o,o
-1 -2 -3 200
250
300
Temp. (°C) Figure 1.190. The relationship between aAg/aAu of native gold and electrum and temperatures. A, aH2s = 10 2, mcl_ = 0.2, inK+ = 0.04, aAuC]£/aAu(HS)£= 102"°, K-feIdspar/K-mica/quartz equlibrium; B ' aH~S = 10 - 2 , m c i - = 0.2 , mK+ = 0.04 ' aAucI /aAu~HSa = 10 T M , C , aH-S = 10 -3 ~ mcl = 0.2 , z ~ /2 z mK+ = 0 . 0 4 , aauCl~-/ a Au(HS)~ = 10 2.0., D, a H 2 S = 10 -~ , me1 = 0 . 2 , mK+ = 0 . 0 4 , aAuCl~/aAu(HS)2 = 10 1.5., E, aAuCI~/aAu(HS)~= 1020; F, aAuci2/aAu(HS)~= 101"5. Calculations were m a d e a s s u m i n g K - f e l d s p a r / K m i c a / q u a r t z equilibrium for A, B, C, and D. T h e r m o c h e m i c a l data for the calculations were taken f r o m Helgeson (1969), Seward (1973, 1976), D r u m m o n d (1981), a n d Henley et al. (1984). (Shikazono a n d Shimizu, 1987).
From the equilibrium relation for equation (1-88) we obtain
aAg/aAu ~ ( a A g c 1 2 K1-88) /aAucI 2 (]-89) Curves E and F in Fig. 1.190 were drawn assuming that the value of N A g / E A u (NAg, total dissolved Ag concentration, NAu, total dissolved Au concentration) is approximately 102-1015 , which corresponds to the ratio for crustal rocks. These curves indicate that the aAg/aAu ratio of native gold or electrum does not change with temperature at a constant E A u / N A g ratio. It is noteworthy that the temperature dependency of the aAg/aAu ratio of native gold and electrum in AuC12 dominant ore fluids (Fig. l. 190). This difference may account for the range of variation and the Ag/Au ratios of native gold and electrum from gold-quartz vein and epithermal A u - A g vein deposits. However, it has to be noted that uncertainties of thermochemical data on gold chloride species might be large. As suggested by Henley et al. (1984) and Brown (1986), it is possible that Ag(HS)2 is dominant in low-salinity and near neutral ore fluids. Under this condition the Ag/Au ratios of electrum and native gold are controlled by the following reaction: (Au) + Ag(HS)~ = (Ag) + Au(HS)2
(1-90)
It is expected from the equilibrium ratios for equation (1-90) that the Ag/Au ratios of native gold and electrum are controlled by temperature and E A g / Z A u in ore fluids, although the equilibrium constant for equation (t-90) has not experimentally been determined. The E A u / E A g ratio in ore fluids responsible for epithermal A u - A g veins could be calculated from equation (1-87). The temperature of formation is taken from fluid
Chapter 1
256 N I
Fukushirr~ K.oganesawa
-~.~Oshima Tochigi Pref.~.'~'--. . ~ ~
Pref.
rune
om, o
Nokoz~mo...{/2/~."
Nosu~ ~ ' . " .... ~_-~z~"'o
• / ' . - Y.,
[: "("Ibaragi Pref. . .) H .) i--Tarozawa Tokakum~... - . . ~ t~etsukuji Otori_.===:/V~N..t---'/:~-'.-. . ." ~--'~Hmaga BATO,~J - .'X.) • xx[ . ,'..qh-.~Shiozawa ~.."-'~.;y /..~'_-_-_-_-_-_-_-.~Kuryu
a UTSUNOMIYA
Figure 1.I91. Locationof gold vein-typedeposits in the YamizoMountains. Dotted area represents the Yamizo Mountains except the Toriashi and Tsukuba blocks. Open circle: Au vein-typedeposits, solid circle: studied Au vein-typedeposits, square: city or town (Shimizu and Shikazono, 1987). inclusion studies (Enjoji and Takenouchi, 1976; Shikazono, 1985b) and the electrumsphalerite-argentite-pyrite assemblage (Shikazono, 1985d). The NaCI equivalent concentration of ore fluids is approximated from freezing data on inclusion fluids (Enjoji and Takenouchi, 1976), though the final melting temperature of fluid inclusion ice is also affected by CO2 concentration in epithermal ore fluids (Hedenquist and Henley, 1985). The pH values are estimated assuming the equilibrium among K-feldspar, K-mica, and quartz, which occur commonly in these deposits; this in turn allows a calculation of the activity of K +. The activity of HzS is estimated based on the equation showing the relation between the partial pressure of H2S gas and temperature for active geothermal waters (Giggenbach, 1980; Arn6rsson, 1985). Using the typical values of these variables and XAg = 0.5, which is a typical value for electrum from Au-Ag veins, E A u / E A g is calculated to be about 10 - I , which is similar to that of Broadlands geothermal water (Table 1.27). The above calculation is based on the assumption that AgCI~- is the predominant Ag species in ore fluids. However, Brown (1986) and Henley (1985) suggested that silver bisulfide complex (Ag(HS)2) could contribute significantly to the transportation of Ag in low-salinity geothermal waters (e.g., Broadlands, New Zealand). Thus, Ag bisulfide complex is also probable as dominant Ag species in ore fluids responsible for Au-Ag veins in Japan, though the salinity of ore fluids (0.1-0.3 molal) is generally higher than that of Broadlands geothermal water. If Au and Ag bisulfide complexes are the dominant Au and Ag species in Broadlands geothermal water and ore fluids responsible for Japanese gold silver veins, the Ag/Au ratio of ore fluids responsible for Japanese
Miocene-Pliocene Hydrothermal Ore Deposits
257
TABLE 1.27 EAu/NAg in ore fluidsestimated from chemical compositionof nativegold and electrum in active geothermal waters and in crustal rocks. XAg (X: atomic fraction) = 0.2 for auriferous veins and XAg = 0.5 for gold-silver veins are assumed. Assumed values include: (1) 200°C, aH2S 10-35, mcl 0.1, pH = 5.6; (2) 250°C, aHzS = 10 2, mcI =0.1, pH = 5.4; (3) 300°C, aH2s = 10-1, mcl- =0.i, pH = 5.4; (4) 300°C; (5) 300°C =
=
Au/Ag log (atomic ratio) (Au/Ag)
References
0.03~0.01 0.013~0.004
-1.6-2.0 -1.1-2.2
Wehdepohl(1978) Holland(1978)
0.i0 0.11 0.03
-0.99 - 1.00 - 1.6
Brown (1986) Henley et al. (1984) Koga (1957, 1961)
0.09
-1.0
(2)
0.09
- 1.0
(3)
0.08
- 1.1
Thermochemicaldata from Helgeson (1969), Seward (1973, 1976), Henley et aI. (1984) Thermochemical data from Helgeson (1969), Seward (1973, 1976), Henley et aI. (1984) Thermochemical data from Helgeson (1969), Seward (1973, 1976), Henley et al. (I984)
Auriferous vein (4)
0.007
--2.2
(5)
1.48 × 10 7
-6.8
(6)
0.007
-2.2
Crustal rocks (average) Seawater Active geothermal water
Broadlands (BR 22) Imperial Valley Beppu Ore fluids
Gold-silver vein (1)
Thermochemical data from Drummond(1981), Seward (I976) Thermochemical data from Helgeson (1969), Seward (1976) Thermochemical data from Drummond(198I), Seward (1976)
gold-silver veins may be similar to that of Broadlands geothermal water which is about 0.1 (Table 1.27). If AgCI~ and AuC12 are the predominant Ag and Au species, we can calculate E A u / E A g in ore fluids by using thermochemical data on gold chloride complexes by D r u m m o n d (1981), Helgeson (1969), and thermochemical mixing properties of A u - A g alloy by White et al. (1957)• The ratio of XAg/XAu (X: atomic fraction) for native gold or electrum from gold-quartz vein deposits is taken to be 0.25. The calculated value of m AuCl£/mAgCl2 based on thermochemical data on AuC12 by D r u m m o n d (1981) is about 10 -2. Using thermochemical data for Au chloride complexes compiled by Helgeson (1969), we obtain mAuci /mA_C1 which is very different from the A u / A g ratios of • 2 g 2 crustal rock and active geothermal waters (Broadlands, New Zealand; Imperial Valley Magmamax 1, California; Beppu, Japan), though the A u / A g values for the Imperial Valley Magmamax 1 and Beppu geothermal water do not reflect those in the deep fluid. ~348 of sulfide (pyrite) from this type of deposit varies widely, ranging from - 5 to +7%0 (Shikazono, unpublished), but average ~34S value is 0%0, suggesting igneous origin from deep part, but some from reduction of seawater sulfate in sediments. 313C values of
258
Chapter i
carbonates from these deposits are mostly from -5%0 to -7%0, suggesting igneous origin (Shikazono, unpublished).
1.6.3.2. Gold-quartz vein-type deposits in Yamizo Mountains, central Japan Shimizu and Shikazono (1987) studied gold-quartz vein-type deposits in Yamizo mountain, central Japan. Their study is summarized below. The Yamizo Mountains extend over the eastern part of Tochigi Prefecture, and the western part of Ibaragi Prefecture, central Japan, occupying an area ca. 20 km by 50 km. A lot of hypo/mesothermal (in the sense of Lindgren, 1933) gold vein-type deposits exist there (Fig. 1.191). The Yamizo Mountains are mostly occupied by Paleozoic-Mesozoic sedimentary rocks, mainly of Jurassic age (e.g., Sashida et al., 1982) with a small amount of intrusive granitoids of unknown ages. The Paleozoic-Mesozoic sedimentary rocks have been called the Yamizo Formation (Kanomata, 1961). It is composed chiefly of shale, sandstone, alternating beds of shale and sandstone, and a small amount of limestone and chert. The succession of the geologic units and geologic structure of the Yamizo Formation have been left pending due to complex structure such as upturned beds (Kasai, 1978) and submarine land sliding (Aono et al., 1985). Many small gold-quartz veins occur in the Yamizo mountains (Fig. I. 191). Some of them were already undertaken under the Satake clan in the Tokugawa era (more than 150 years ago). Most of them were productive during the 1930s to 1950s in several ten tons per month with a grade of ten to 50 grams per metric ton of gold and silver, respectively. Gold-quartz veins ranging in width from several centimeters to several ten centimeters are situated in shale and/or sandstone of the Yamizo Formation and rarely in granitoids. The strike and dip of the veins are various. For example, the veins of the Kuryu, Shiozawa, and Daigo deposits strike N40°W, N20-30°W, and N30°E, respectively, and dip 60°E, 30-50°W, and 20°W, respectively. The veins are composed mostly of quartz and a small amount of sulfide minerals (pyrite, pyrrhotite, arsenopyrite, chalcopyrite, sphalerite, and galena), carbonate minerals (calcite, dolomite) and gold, and include breccias of the host rocks with carbonaceous matters. Layering by carbonaceous matters has been occasionally observed in the veins. Banding structure, wall rock alteration and an evidence of boiling of fluids that are commonly observed in epithermal veins have not been usually found. Gold is present either as isolated grains in quartz or in association with sulfide minerals. In the case of the latter, gold is generally in direct contact with pyrite and/or galena. Gold is irregular in shape with 0.01 to l mm in size. Some of gold grains can often be observed by naked eyes. Analytical results on gold are shown in Table 1.28. Chemical compositions of gold are relatively homogeneous and low Ag contents. Frequency histograms for Ag atomic% of gold (Nag) from the Yamizo Mountains and other regions are summarized in Fig. 1.192. Values of NAg of gold from the Yamizo Mountains range from 5.9 to 18.4, and this is almost the same as those from the other regions (Fig. 1.192). However, NAg of hypo/mesothermal gold vein-type deposits in metamorphic rocks tends to be slightly lower than that in sedimentary rocks.
Miocene-Pliocene Hydrothermal Ore Deposits
259
TABLE 1.28 Representative chemical compositions of gold from the Yamizo Mountains, Japan (Shimizu and Shikazono, 1987) Ore deposit
Weight % Au
Atomic % Ag
Cu
Total
Au
Ag
Cu
Ag/Au
Kuryu
96.70 96.31 96.15 96.13 96.20 95.66 95.89 96.05 95.36 95.24
3.31 3.41 3.50 3.6I 3.77 3.79 3.90 3.97 4.12 4.26
0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00
100.01 99.72 99.65 99.74 99.97 99.46 99.79 100.02 99.48 99.50
94.11 93.93 93.78 93.58 93.33 93.22 93. l0 92.98 92.69 92.45
5.89 6.07 6.22 6.42 6.67 6.74 6.90 7.02 7.31 7.55
0.00 0.00 0.00 0.00 0.00 0.04 0.00 0.00 0.00 0.00
0.063 0.065 0.067 0.069 0.071 0.072 0.074 0.076 0.079 0.082
Shiozawa
92.83 92.10 92.24 92.15 92.22 92.62 92.20 91.96 91.55 91.64
7.00 7.06 7.11 7.16 7.24 7.37 7.37 7.45 7.48 7.55
0.05 0.05 0.02 0.04 0.06 0.03. 0.06 0.06 0.04 0.00
99.88 99.21 99.37 99.35 99.52 100.02 99.63 99.47 99.07 99.19
87.76 87.60 87.61 87.47 87.32 87.24 87.12 85.98 86.92 86.92
12.09 12.25 12.33 12.42 12.52 12.67 12.71 12.86 12.96 13.08
0.15 0.15 0.06 0.11 0.17 0.09 0.17 0.17 0.11 0.00
0.138 0.140 0.141 0.142 0.143 0.145 0.146 0.150 0.149 0.150
Daigo
92.32 91.68 91.65 91.77 91.54 91.64 91.27 91.41 91.33 91.16
7.54 7.5I 7.69 7.78 7.83 7.93 8.04 8.11 8.18 8.37
0.02 0.01 0.02 0.02 0.02 0.03 0.02 0.04 0.00 0.04
99.88 99.20 99.36 99.57 99.39 99.60 99.33 99.56 99.51 99.57
86.97 86.96 86.66 86.55 86.44 86.28 86.10 85.96 85.95 85.54
12.97 13.00 13.28 13.40 13.50 13.63 13.85 13.93 14.05 14.35
0.06 0.04 0.06 0.06 0.06 0.09 0.06 0.11 0.00 0.11
0.149 0.149 0.153 0.155 0.156 0.158 0.161 0.162 0.163 0.168
Saigane
89.73 89.40 89.56 88.69 89.19 89.12 88.42 88.99 88.30 88.34
9.68 9.71 10.19 10.28 10.46 10.54 10.66 10.78 10.87 10.91
0.07 0.06 0.06 0.06 0.08 0.08 0.07 0.08 0.08 0.09
99.48 99.17 99.81 99.03 99.73 99.74 99.15 99.25 99.25 99.40
83.38 83.31 82.67 82.39 83.17 82.05 81.79 81.46 81.46 81.31
16.42 16.52 17.17 17.44 17.59 17.72 18.01 I8.30 18.30 18.44
0.20 0.17 0.16 0.16 0.24 0.24 0.20 0.24 0.24 0.25
0.197 0.198 0.208 0.212 0.214 0.216 0.220 0.225 0.225 0.225
Chapter 1
260 N=144
In Metamorphic R. Suwa, Kinkef, Amo, Hashidate
50 4,0 30 20 10
El N: 1[
70
Sh
6O 5O
In Sedimentary R. Yamizo Mtns, Kuryu, Daigo Shiozawa, Saigane
D
o 40 o-
~-
30
u_ 2O
10 ~
~Sa
S. Koma Region Ho Gohaku, Koei
[
N=1'57j~ 30. A~erlgel)
' Kitakami Mtns. Oya, Shishiori,
20
Hayachine, Kohoku Nedachi(1974) Yamaoka(1981)
10
0
10 20
50
100 NAg
Figure 1.192. Frequency (no. of analyses) histograms for Ag content (in atomic %) of gold from "hypo/ mesothermal" gold vein-typedeposits in Japan. K: Kuryu deposit, Sh: Shiozawa deposit, D: Daigo deposit, Sa: Saigane deposit. Data are from: the Yamizo Mountains:Shikazonoand Shimizu (1988a); the South Korea Region: Shikazono and Shimizu (1987); the Kitakami Mountains: Nedachi (1974); Yamaoka (1981), Abe (1981) and Shikazono and Shimizu (1987); metamorphicregions: Shikazono and Shimizu (t987).
In the Yamizo Mountains, the correlation between NAg and a kind of the host rocks is found: NAg of gold from gold-quartz veins in shale is lower than that in sandstone. Gold negatively correlates with 834S of the associated sulfide minerals from the veins, and that 334S of sulfide minerals in sandstone is lower than that in alternating beds of shale and sandstone, probably than that in shale.
Miocene-Pliocene Hydrothermal Ore Deposits
261
Homogenization temperatures of fluid inclusions mainly in quartz closely associated with gold and sulfide minerals in hypo/mesothermal gold vein-type deposits in Japan are in a range from 250 ° to 350°C (Shikazono and Shimizu, 1987), higher than those in epithermal gold-silver vein-type deposits in Japan. Iron contents of sphalerite and thermochemical data by Barton and Toulmin (1966) and Barton and Skinner (1979), possible sulfur fugacity (fs2) of the ore fulids can be estimated: log fs2 = - 1 5 at 250°C to - 3 at 350°C. This means that fs2 for hypo/mesothermal gold vein-type deposits is lower than that for epithermal gold-silver vein-type deposits in Japan. Factors in controlling chemical compositions of gold in equilibrium with the ore fluids are temperature, pH, concentration of aqueous H2S and C1- in the ore fluids, concentration ratio of Au and Ag species in the ore fluids, activity coefficient of Au and Ag components in gold, and so on (Shikazono, 1981). In the Yamizo Mountains, as a result, Ag/Au ratios of gold are correlated with a kind of the host rocks and sulfur isotopic compositions of the deposits. This correlation could be used to interpret Ag/Au ratios of gold. There are two possibilities here to explain this correlation. One is that isotopically heavy sulfide sulfur derived from seawater sulfate was fixed in shale because reducing agency of shale with carbonaceous matters is thought to be stronger than that of sandstone. The ore fluids extracted this sulfur. Gold of low NAB precipitated in shale like the Kuryu deposit under more reducing environment than in sandstone like the Saigane deposit. Shikazono and Shimizu (1987) concluded that Ag contents of gold precipitated from low-salinity fluids is higher than that prediction and the relationship between Nag of gold and salinity of fluid inclusions estimated from freezing temperature data. Therefore, another interpretation is that Nag of gold from shale-hosted deposits is lower than that from sandstone-hosted deposits, because shale is expected to be richer in C1- mainly due to adsorption by clay minerals included in shale than sandstone. It is concluded that, in either case, in situ interaction of the ore fluids with the host rocks controls the chemical compositions of gold in the Yamizo Mountains.
1.6.4. Hot spring-type gold deposits Hot spring-type gold deposits (Nansatsu-type by Urashima et al. (1981, 1987), high sulfidation-type by Hedenquist (1987), epithermal Au disseminated-type) are distributed in the Nansatsu district of southern Kyushu (Fig. 1.193). The deposits (Kasuga, Akeshi, Iwato) were formed at Pliocene age (5.5-3.7 m.y.) in the calc-alkaline volcanic rocks of nearly same age (Togashi and Shibata, 1984). The deposits, which are similar to Nansatsu-type deposits, occur in Southwest Hokkaido (Date, Hakurhu). The deposits are characterized by conspicuous alteration zoning from the centre (orebody) to margin (Tokunaga, 1955; Doi, 1972; Urashima et al., 1981, 1987). They are siliceous zone, alunite zone, kaolinite zone, sericite zone and montmorillonite zone. In the siliceous body, electrum and Cu minerals (enargite, luzonite, covelline), and native sulfur occur. The Ag content of electrum is lower (0.0-5.3 wt%) than that from epithermal A u - A g vein-type deposits (Fig. 1.194) (Shikazono and Shimizu, 1987). Low Ag content of electrum and sulfide mineral assemblage (enargite, native sulfur, covellite, pyrite) indicate high i s 2 condition (Fig. 1.194).
Chapter I
262
J~ f J
~ ~
\ o
~._~v-
~
[ [
- ~
~
• Hot springtype
~ EpithermalAu-Agvein-type ~ Active volcano
o
20
I,
~,
4o
I
/
o o~uo~i
)
/
o,.o..=o
(
Kagoshima
/
j
Mt.Kaimondake
31N
I
~
l 131E
Figure 1.193. Locality map of the hot spring-type deposits in Kyushu. 4O
=~ 3o
g
2O
10
0
0
20
40
60
80
1oo NAg
Figure 1.194. Frequency histogram for the Ag content of eIectrum from epithermal Au disseminated-type deposits in Japan (Iwato and Kasuga). (Shikazono and Shimizu, 1987).
Miocene-Pliocene Hydrothermal Ore Deposits
263
/
Steam-heated alunite fluid / / / I .i~<-~/
•
K ~
I 0 Alunite
•
0 Clays
SMOW
.~'V Alunitefluid '@ ~
{ 1._.: ~s~ ~
volcanic fluids
O
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a
Clayfluid 0 ~ 0 0
-100
.2
I
I
t
I
Residual silica • I
(
I
H
•
silica
Kasuga
and
0
O Iwato
Deepveins at Kasuga
(~
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~
vein
quartz
Other localities Calculated fluids
I
-10
•
0 Residual
Quartz veins in residualsilica
I
0
8180
+10 %o
I
+20
Figure 1.I95. Diagram showing the results of isotopic analysis of residual silica (now recrystaltized to quartz) and quartz crystals or veins (3180 only, plotted on the bottom), and clays and alunite (~D vs. 3180, plotted on the top). The calculated composition of the fluids in equilibrium with these minerals, plus the compositions of local meteoric water (small open square) and high-temperature volcanic fluids (large open square) are also shown (Matsuo et al., 1974; Matsubaya et al., 1975). A least-squares fit line (dashed) through the hypogene alunite fluids and the high-temperature volcanic fluid end member extends to an isotopic composition similar to present local meteoric water and identical in gD values to the clay fluid. Thus the isotopic composition of alunites record mixing of magmatic and meteoric waters, whereas the marginal clays (and possibly the vein quartz and residual silica quartz) indicate that meteoric water dominated on the margins of the quartz bodies, variably shifted in oxygen isotope composition due to exchange with wall rocks. Some of the deep fluids responsible for vein formation are identical in isotopic composition to the 180-enriched fluid that equilibrated with the residual silica quartz, whereas other samples indicate a greater proportion of unshifted meteoric water. One vein alunite from Iwato has an isotopic composition which indicates that it formed in a steam-heated environment after evaporation of local ground water (fluid composition calculated for a steam-heated temperature range of 100-180°C). K and I denote Kasuga and Iwato, respectively. (Hedenquist et al., 1994).
Chapter 1
264
Hedenquist et al. (1994) indicated that homogenization temperature of postmineralization of the Akeshi deposit and salinity of ca. 1 wt% NaCl eq. and the presence at times of a two-phase fluid in the centre of ore zones indicating boiling of ore fluids. 834S of sulfides (pyrite, enargite, luzonite) and sulfate(alunite) are -5%0 to 0%0 and +24%0 to +35%0, respectively (Shikazono, 1987b; Hedenquist et al., 1994), suggesting that sulfide sulfur came both from igneous and sedimentary sources and sulfate sulfur is generated by the hydrolysis reaction of 4 S O 2 -k- 4H20 ~ H2S + 3 H 2 8 0 4 . SO2 gas may come from volcanic gas. 3D and 8180 of ore fluids were estimated from these data on quartz, clay minerals and alunite (Fig. 1.195) (Hedenquist et al., 1994), indicating higher 8D and 8180 values compared with epithermal Au-Ag vein ore fluids. These isotopic data (834S, 8D, 8180), fluid inclusion data and alteration studies suggest that (1) the gas-rich (HCl, SO2, H2S) vapor phase separated from magma ascent to the surface and condense into meteoric water, (2) this acid chloride-sulfate fluid leached elements from the host rocks, forming a permeable zone of mixing, (3) then, the dense metal-rich liquids ascent into the leached zone and precipitate Cu-minerals and Au (Hedenquist et al., 1994). The fo2-PH conditions for this type of deposits and other types of epithermal deposits and hydrothermal alteration are shown in Fig. 1.196 (Shikazono and Aoki, 1981). It is worth noting that the f o 2 - P H ranges for the epithermal Au deposits (hot spring-type, Te-type, and Se-type) lie between the points A (high fo2, and low pH) and B (low f o e
0
Log fo 2 'o~ vm, e High sulfidation type
'r ,
I
I
/ll
~AIIKal , / 1 , Y"
-30 •
."--. ,,
l./I
Low sulfidation type
/I /'
/ s
s,
-
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Bn+Py Cp
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~
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--I---I
~40
i
I
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I
I
2
3
4
I
5
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'
4
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1
e4~CO
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~
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I
Se I
~.
' ""!"';"c'--
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1
I
I~
|
I
6
7
8
9
pH
Figure 1.196. fo2 pH ranges for hot-spring-type deposits and low sulfidation-type deposits. Temperature = 250°C, £S = 0.01 mol/kg H20, ionic strength = I. Ka: kaolinite, AI: alunite, SI: liquid sulfur, Kf: K-feldspar, Hm: hematite, Mt: magnetite, Py: pyrite, Po: pyrrhotite, Bn: bornite, Cp: chalcopyrite.
Miocene-Pliocene Hydrothermal Ore Deposits Assumed direction of the Pacific Plate Deviation form the north 90 ° 70 ° 50o 30 °
Ma 0
EpithermaI system
Q.R intrusion
~
O
12 -
f
N
3.
- Tokiwamatsu H.S.S. - Kobui H.S.S.
.
4 5-
265
-O
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Hokko-Minami LS.S.
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6--Mitsumoriyama
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Q.R
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8N 910-
H.S.S.: High sutfidation s y s t e m L.S.S.: L o w sulfidation s y s t e m Q.R: Q u a r t z p o r p h y r y N: n o r m a l s u b d u c t i o n , O: o b l i q u e s u b d u c t i o n Figure 1.197. Subduction mode of the Pacific plate beneath the Northeast Japan arc, and style of the epithermal systems in the Kameda peninsula, Hokkaido. Normal, oblique and highly oblique subduction are tentatively defined by the angle (c0 between the assumed subduction direction of the Pacific plate and the trend of the Northeast Japan arc. Normal subduction (60 ° < c~ = 90°), oblique subduction (30 ° < c~ = 60 °) and highly oblique subduction (0° < c~ = 30°). Highly oblique subduction has not occurred along the Northeast Japan arc since 11 Ma. (Watanabe et al., I996).
and high pH) in Fig. 1.196. These difference in physicochemical nature of fluids can be explained by space-dependent evolution of a hydrothermal system relative to the centre of volcanic activity (Shikazono and Aoki, 1981; Shikazono, 1985a; Shikazono et al., 1990; Sillitoe, 1991) or a time-dependent evolution of a hydrothermal system (early-stage; high sulfidation-type, late-stage; low sulfidation-type) (Aoki et al., 1993). For example, Shikazono (1985a) has shown that advanced argillic alteration (Ugusu silica deposit in west Izu Penninsula, central Honshu) occurs at the centre and
Chapter 1
266
stratigraphically higher position in the mine district, while epithermal Au-Ag veintype deposits at the margin and lower position (Fig. 1.76). Shikazono et al. (1990) compared the characteristic features of Te-type with Se-type and suggested that the Te-type formed at the position closer to volcanic centre in a hydrothermal system than Se-type (Fig. 1.123). It is probable that the difference in fo2 and pH in several types of epithermal Au deposits is due to the different ratio of meteoric water component (low fo2 and high pH) to magmatic water (or igneous water) component (high fo2, low pH), related to the position from volcanic centre. Watanabe et al. (1996) and Watanabe and Ohta (1999) pointed out that the high sulfidation (advanced argillic alteration) and low sulfidation (sericitic and propylitic alteration) system in Southwest Hokkaido (Minamikayabe area, Jozankei-Zenibako district) occurred during normal and oblique subduction of the Pacific plate beneath the Northeast Japan arc, respectively (Fig. 1.197). These studies suggest that the difference in two styles of epithermal gold mineralization (epithermal vein (low sulfidation-type) and hot spring-type (high sulfidation-type)) is caused by the subduction mode of the plate and not by the temporal and spatial differences in a given hydrothermal system. Figure 1.198 shows grade-tonnage relationship for low sulfidation-type (Se-type) and high sulfidation-type (Te-type and Akeshi-type) in Japan and deposits in Circum Pacific region summarized by Hedenquist et al. (1996). This clearly indicates that grade and tonnages for Akeshi-type and Te-type in Japan are low and small, compared with low sulfidation-type in Japan. It is obvious that low sulfidation-type is dominant Au-Ag deposits in Japan. The plot of K20 versus SiO2 for volcanic rocks thought to be genetically related to Au mineralization indicates that high sulfidation deposits (hot spring-type deposits appear
ll-t
100
I
Illllll
I Xl~
"'.z°o, -. 7 "za
"
-.e %
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]0
o", ",
eO OO
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I
I
~
I
[[IFF[
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I
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%0
,.L.,
. ' ~ "" ~
•
•
""
" "•-o RO • • "~'oEmMc oIB~'" "PV La "e. • • p p
] 0.l
[
"" f.ooa "- z Hi "'.
1
10
,e. A 100
RM ". 1000
Tonnage, Mt Figure 1.198. Grade-tonnage plot for epithermal Au deposits. LS: Ba, Baguio (Philippines); CC, Cripple Creek; Mc, McLaughlin and RM, Round Mountain (USA); Era, Emperor (Fiji); Hi, Hishikari (Japan); Ke, Kelian (Indonesia); La, Ladolam and Po, Porgera Zone VII (Papua New Guinea). HS: Ch, Chinkuashih (Taiwan); EI, EI Indio (Chile); Go, Gotdfieid and PP, Paradise Peak (USA); Le, Lepanto (Philippines); PV, Pueblo Viejo (DominicanRepublic); Ro, Rodalquilar (Spain) (Hedenquist et al., 1996).
Miocene-Pliocene Hydrothermal Ore Deposits 8
i
J
i
267
i
i
t
[
• LS(calc-alkalic) • LS(shoshonitic-alkalic) • HS(caIc-alkalic)
6
jO 4~
4
4 ~
•~
• ~0
•
._
•
L-A
•
•
•
• High K
~
Medium-K
¢, •
I
O0
Low-K I
50
I
l
60 SiO2,wt%
I
I
70
I
80
Figure 1.199. Plot of K20 versus SiO2 for volcanic rocks thought to be genetically related to Au mineralization. High sulfidation (HS) deposits appear to be associated with a narrow range of igneous rock composition, dominated by dacite and andesite. Low sulfidation (LS) deposits are largely associated with andesite and rhyolite, but major deposits such as Ladolam, Porgera and Cripple Creek are associated with shoshonitic and alkalic rocks (Hedenquist et al., 1996).
to b e a s s o c i a t e d w i t h a n a r r o w r a n g e o f i g n e o u s r o c k c o m p o s i t i o n , d o m i n a t e d b y d a c i t e a n d a n d e s i t e , w h i l e l o w s u l f i d a t i o n d e p o s i t s are w i t h a n d e s i t e a n d r h y o l i t e , b u t m a j o r d e p o s i t s (e.g., L a d o l d m , P o r g e r a a n d C r i p p l e C r e c k ) are a s s o c i a t e d w i t h s h o s h o n i t i c a n d alkalic r o c k s ( H e d e n q u i s t et al., 1996) (Fig. 1.199).
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Bischoff, J.L. and Dickson, EW. (1975) Seawater-basalt interaction at 200°C and 500 bars; Implications for origin of seafloor heavy-metal deposits and regulation of seawater chemistry. Earth Planet. Sci. Lett., 25, 385-397. Bischoff, J.L., Radke, A.S. and Rosenbauer, R. (1981) HydrothermaI alteration of graywacke by brine and seawater: rocks of alteration and chloride complexing on metal solubilization at 200°C and 350°C. Econ. Geol., 76, 659-676. Blount, C.W. (1974) Synthesis of barite, celestite, anglesite, witherite, and strontianite from aqueous solutions. Am. Min., 59, I209-1219. Blount, C.W. (1977) Barite solubilities and thermodynamic quantities up to 300°C and 1400 bars. Am. Mineral., 62, 942-957. Blow, W.H. (1969) Late middle Eocene to Recent planktonic foraminiferal biostratigraphy: 1st International Conference on Planktonic microfossils, Proceeding, 1, 199-421. Bonham, Jr., H.E (1984) Three major types of epithermal precious-metal deposits (abst.). Geol. Soc. Am. Abst. with Programs, 16, 449. Bonham, Jr., H.E (1986) Models for volcanic-hosted epithermal precious metal deposits; a review. Internat. VolcanoI. Congr., Hamilton, New Zealand, Feb. 1986, Symposium V. Proc., 13 17. Boorman, R.S. (1967) Subsolidus studies in the ZnS-FeS-FeS2 system. Econ. Geol., 62, 614-631. Bowers, T.S. and Taylor, H.R Jr. (1985) An integrated chemical and stable isotope model of the origin of midocean ridge hot spring systems. J. Geophys. Res., 90, 12583-12606. Bowers, T.S., Von Damm, K. and Edmond, J.M. (1985) Chemical evolution of midocean ridge hot springs. Geochim. Cosmochim. Acta, 49, 2239-2252. Brown, K.L. (1986) Gold deposition from geothermal discharges in New Zealand. Econ. Geol., 81, 979-986. Browne, RR.L. and Ellis, A.J. (1970) The Ohaki-Broadlands hydrothermal area, New Zealand: Mineralogy and related geochemistry. Am. J. Sci., 269, 97 131. Browne, RR.L. (1978) HydrothermaI alteration in active geothermal fluids. Ann. Rev. Earth Planet. Sci., 6, 229-250. Bryndzia, L.T., Scott, S.D. and Farr, J.E. (1983) Mineralogy, geochemistry, and mineral chemistry of siliceous ore and altered footwall rocks in the Uwamuki 2 and 4 deposits, Kosaka mine, Hokuroku district, Japan. Econ. Geol. Mon., 5, 507-522. Buchanan, L.J. (1981) Precious metal deposits associated with volcanic environments in the southwest. In: Relations of Tectonics to Ore Deposits in the Southern Cordillera (Dickson, W.R. and Payne, W.D., eds.), Arizona Geological Society Digest, 14, 237-262. Buddington, A.E (1935) High-temperature mineral association at shallow to moderate depth. Econ. Geol., 31, 205-222. Burnham, C.W. and Ohmoto, H. (1980) Late-stage process of felsic magmatism. Soc. Mining Geologist. Japan, Special Issue, 8, 1-11. Cart, M.J., Stoiber, R.E. and Drake, C.L. (1973) Discontinuities in the deep seismic zones under the Japanese arcs. Geol. Soc. Am. BulL, 84, 2917-2983. Casadevall, T. and Ohmoto, H. (I977) Sunnyside mine, Eureka mining district, Sun Juan County; Geochemistry of gold and base-metal ore deposition in a volcanic environment. Econ. Geol., 72, i285-1320. Cathelineau, M. and Nieva, D. (1985) A chlorite geothermometer. The Los Azufres (Mexico) geothermal system. Contr. Mineral. Petrol., 91,235-244. Cathles, L.M. (1983a) Kuroko-type massive sulfide deposits of Japan: Products of an aborted Island-arc rift. Econ. Geol. Mon., 5, 96-114. Cathles, L.M. (1983b) An analysis of the hydrothermaI system responsible for massive sulfide deposition in the Hokuroku basin of Japan. Econ. Geol. Mort., 5, 439-487. Cathles, L.M. (1986) The geologic solubiIity of gold from 200-300°C, and its implications for gold-base metal ratios in vein and strafiform deposits. In: Clark, L. (ed.), Gold in Western Shield. Can. Inst. Min. Metall. Spec., 38, 187-210. Cavaretta, G., Gianelli, G. and Puxeddu, M. (1980) Hydrothennal metamorphism in the Larderello geothermal field. Geothermics, 9, 297-314. Cavaretta, G., Gianelli, G. and Puxedolu, M. (1982) Formation of authigenic minerals and their use as
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