Chapter 3 Hydrothermal flux from back-arc basin and island arc and global geochemical cycle

407

Chapter 3

Hydrothermal Flux from Back-Arc Basin and Island Arc and Global Geochemical Cycle

3.1. Major element (alkali, alkali earth, silica) flux It was shown in previous chapters that intense hydrothermal activities occurred in the Neogene age in and around the Japanese Islands under the submarine and subaerial environments. In this chapter the influence of these hydrothermal activities on the seawater chemistry, and the global geochemical cycle are considered. The studies on the hydrothermal systems at midoceanic ridges during the last three decades clearly revealed that the seawater-basalt interaction at elevated temperatures (ca. 100-400~ affects the present-day seawater chemistry (Wolery and Sleep, 1976; Edmond et al., 1979; Humphris and Thompson, 1978). For example, a large quantity of Mg in seawater is taken from seawater interacting with midoceanic ridge basalt, whereas Ca, K, Rb, Li, Ba and Si are leached from basalt and are removed to seawater (Edmond et al., 1979; Von Damm et al., 1985a,b). As mentioned already in Chapter 2, submarine volcanism occurs not only at midoceanic ridges but also at subduction-related tectonic settings such as the Shikoku and Daito Basins, Parce Vela Basins, and Mariana Trough, Okinawa Trough and Izu Bonin Arc (e.g., Wood et al., 1980; Dick, 1982; Delaney and Boyle, 1986). We saw in section 2.3.2 that present-day hot spring venting and sulfide-sulfate depositions have been discovered in back-arc basins in the Western Pacific. These intense hydrothermal activities indicate that seawater-volcanic rock interactions are taking place at these environments. Bulk rock chemistry of hydrothermally altered midoceanic ridge basalt has been well studied and used to estimate the geochemical mass balances of oceans today (Wolery and Sleep, 1976; Humphris and Thompson, 1978; Mottl, 1983). In contrast, very few analytical data on hydrothermally altered volcanic rocks that recently erupted at back-arc basins are available. However, a large number of analytical data have been accumulated on the hydrothermally altered Miocene volcanic rocks from the Green tuff region in the Japanese Islands which are inferred to have erupted in a back-arc tectonic setting (section 1.5.3). The age of Green tuff volcanic activity ranges widely from ca. 25 Ma to 2 Ma. Volcanic activity during the early to middle Miocene (25-15 Ma) was intensive, whereas it was weak during the late Miocene to early Quaternary (Sugimura et al., 1963) (Fig. 3.1). The production of lavas and other effusives per unit time reached five or six time more

Chapter 3

408

Early Neogene ( 150.000 km3) 9

~

t

~

i

~176

Acidic Rock ~] Basic Rock

(10.000 km 3) Quaternary

J .000 Middle and Late Neogene 20.000 e.~

0

5

9

e.o

o

10

15

20

25 (M.Y)

Fig. 3.1. Histogram showing the change in volume of volcanic products during the Neogene and Quaternary periods. The far-right part shows a very small amount for the Paleogene Period (Sugimura et al., 1963). than that of the Quaternary period inclusive of the present time (Ozawa, 1963; Sugimura and Uyeda, 1973). Recent K-Ar age data suggest that the volcanic activity during the Nishikurosawa stage (16-14 Ma) was the most intense one (Tsunakawa and Takeuchi, 1986). The volcanic activity during this stage was bimodal, characterized by basaltic and dacitic activities (Konda, 1974). Many drillings have been made for the purpose of the exploration of Kuroko deposits in the Green tuff region. A large number of petrographical and analytical data are available. Thus, it is possible to (1) assess the proportion of volcanic rocks that have suffered hydrothermal alteration, and (2) attempt a quantitative evaluation of the influence of hydrothermal alterations at back-arc basins on middle Miocene seawater chemistry. The Mg content of hydrothermally altered volcanic rocks is reflected by the extent of seawater-volcanic rock interaction at elevated temperatures, because it has been experimentally and thermodynamically determined that nearly all of the Mg in seawater transfer to volcanic rocks, owing to the reaction of the cycled seawater with volcanic rocks at elevated temperatures (Bischoff and Dickson, 1975; Mottl and Holland, 1978; Wolery, 1979; Hajash and Chandler, 1981; Reed, 1983; Seyfried, 1987). It has been shown that the CaO content of hydrothermally altered midoceanic ridge basalt is inversely correlated with the MgO content with a slope of approximately - 1 on a molar basis (Mottl, 1983). This indicates that Ca of basalt is removed to seawater and Mg is taken up from seawater by the formation of chlorite and smectite during the seawater-basalt interaction. This type of reaction is simply written as: CaO + Mg 2+ --+ MgO + Ca 2+ basalt seawater chlorite, smectite modified seawater

(3-1)

A large number of analytical data on the altered volcanic rocks from Green tuff regions in Japan are summarized in Fig. 3.2. Although data are scattered, the data

Hydrothermal Flux from Back-Arc Basin and Island Arc

(a)

12

12 "4 J

'Je

409

(b)

9

9 o

9e

10 P

9

9 ee i

eo

9

,

10

9

e

9 0

8 9 9

o4 v

9

e

9

eeoo

9 .

o

e

9

9

eee

.,~;,.

.

Oe ,eO 9

6

coo oo

0 0

9

9

9

e9 9

v

.% 9

" o0

9

4~e

9

4

.

dJ

9 9 co.

f

,~. IMP

9 9

Je

9

, , .

oj

g"

9

eee

9

9 9

.5

oil

I 8'

e 9

2

o#

9

o e

9 9

|

'~ 9

i

0 0

2

4

i

6

6

e 9

e~

9

.*

"0

9

,'.

Je

i 9 o,

9

-S ew

Q9

,,;'i " ~ 9 ~r 2,~.-4,.-

9

,

I' 9

.p~,,

9

2

4

s

i

i

t

8

10

12

MgO (wt %)

81

14

l

16

0

I

0

6

f

8

J

10 12

14

16

MgO (wt %)

Fig. 3.2. (a) The relationship between MgO and CaO contents of basaltic rocks and of (b) dacitic rocks from the Green tuff region in Japan (Shikazono, 1994).

demonstrate that the Mg in seawater exchanged for Ca of basalt on a molar basis. This relationship between the MgO and CaO contents of the basalt from Green tuff region is similar to that for the hydrothermally altered midoceanic ridge basalt (Mottl, 1983). The relationships between MgO content and the contents of the other constituents (K20, SiO2) of the hydrothermally altered volcanic rocks from Green tuff region are also similar to those for the hydrothermally altered midoceanic ridge basalt. It is indicated that SiO2 and K20 contents of basalt and dacite are taken up by the cycled seawater. From the difference between the MgO content of fresh and altered volcanic rocks, Mg uptake by volcanic rocks is estimated to be ca. 1-10 g per 100 g basalt and 1-5 g per 100 g dacitic rocks. It is notable that the MgO uptake by basalt from Green tuff region is similar to that by midoceanic ridge basalt at ridge axis, but that by dacitic rocks is smaller. Gain and loss of other elements (Ca, K, Si) can be also estimated based on the value of Mg uptake and the relationship between MgO content and the contents of these elements. The total volume of volcanic rocks that erupted during the Green tuff volcanic activity (25-2 Ma) is estimated to be 617,000 km 3 (Sugimura et al., 1963). During the early to middle Miocene age (25-15 Ma) the volume was large (15,000 km3), whereas it was small (2,000 km 3) during the late Miocene to Pliocene (Fig. 3.1) (Sugimura et al., 1963). The total volume ratio of acidic to basic volcanic rocks that erupted in the early to middle Miocene is 3 : 2 (Sugimura et al., 1963). Therefore, the most reasonable average Mg uptake is 2-3 g/100 g volcanic rocks. Using this value and duration of volcanic

410

Chapter 3

activity in the early to middle Miocene (about 10 million years), the average rate of annual Mg removal from seawater to volcanic rocks during 25-15 Ma is estimated to be 1 • 0.2 x 1012 g/year. The total mass of volcanic rocks that erupted per year during 25-15 Ma is estimated to be 4 • 1013 g/year. The rate of seawater cycling is 7.7 4- 1.5 • 1014 g/year, if all of the Mg in cycled seawater is removed to volcanic rocks by the reaction of cycled seawater to rocks at elevated temperatures. Thus the average seawater/volcanic rock ratio (by weight) is calculated to be 14 4- 7 which is roughly similar to that estimated for the recharge zone of the present-day midoceanic ridge hydrothermal system (e.g., 10 + 8 by weight; Humphris and Thompson, 1978). Intense submarine and subaerial volcanic activities during the Tertiary at Green tuff regions took place not only at the Japan Sea but also at marginal basins in the circum-Pacific Region. According to the summary of the development of back-arc basins in the Cenozoic age by Tamaki and Honza (1991) (Figs. 3.3 and Fig. 3.4) and Kaiho and Saito (1994) (Fig. 3.5), many back-arc basins (Japan Sea, Kuri, Shikoku, Parece Vela, South China, Sulu, Makassar, Central Scotia, Cayman) widely and rapidly developed during 30-15 Ma. The total volume of volcanic rocks that erupted at back-arc basins in the circumPacific region during the early to middle Miocene is difficult to estimate. However, it is likely that the total volume of submarine volcanic rocks that erupted during early to 65 ~

,/ Komandosky Aleutian Q

,. C a y m a n

Caribbean Ib

;ce Vela Carohne Manus

~ 9 Loyalty

9'

'W

#

9

Lau 1.4avre

d ~

#

tqP'~

New Caled. . . .

m~.-]

Emerald IO

,.,y,

65~

180~

20~

Fig. 3.3. Marginal basins around the Pacific Ocean. Solid areas show marginal basins, and blocked lines show trenches. E: East, N: North, S: South, W: West (Tamaki and Honza, 1991).

Hydrothermal Flux from Back-Arc Basin and Island Arc BASIN~

E

O ;;" i

Aleutian Komandosky Kudl Japan Shikoku Oklnawa

~

Madana

0 z

Parece Vela W. Philippine South China Sulu Celebes Banda Molucca Makassar Andaman Caroline Manus Solomon Woodlark Coral Loyalty North Fiji Lau-Havre South i%ji New Caledonia Emerald Tasman

Central Scotia East Scotia m Caribbean <: Cayman

I OLIGOCENE I.

IQIPL'I

20 t

Oi Ma

I II

I

iI II

30 t

I I

II ,==

i!

I

,,,,

i I

I I I! I

I

I

I

I I I ,._.. I I I I |

! 1

I

<

I

tI I iII

I I I

I

I I I I II II

I

-o.f I II

I

I II

I 1

i !

I!

I( II I I I I I II I I I I|

,

I

I

I i

I I ! I I

I I

J

f I| I

L!

I I !

I

I I II I

i II

I

I I

I II

I! I! II 1

i, II

I

1I I I I

I I I I

I

1 IIII I~lillUl

I! II I! 1I

1 II ili

Illllll~

I

~11II IIl II II hill l I IIII I

1 I

I

I 1 I! II

Illlllllll

i I

I II

I I I

I

!

I I

I

III II n

I! 1 I ! I !

I ! !

I

1t

I I

II

I

I|

| )

II

II II I

II I Ii I1 I I II

I

I II I |

I

I I II II

I

I II

I

I!

I II

I

I II II I I II I

1 I!

i I

! I

I

I I

II

I 1 1I

It

I

II

!

I ! ! ! !

I I

II

1! I!

I 1i II II I! I I II I 1 1!

~0

II

I I

I II

7O i

I I I II II

I

II

CRETACEOUS

!

II I ( I II I I

I PALEOCENEI

.op

!

', I I

!

II II I I II I

EOCENE

I I I i

Illllilkll

I

I

4,0

411

II I II

I

I I f lilllilliil

I I

I I I If I I I

I I

I

! i ! t

1

! 1

1 1 1 I I I

1!

I I I

1I 1

1I

I II 1I I

II II I

I

I! I 1 1

Fig. 3.4. Range chart of back-arc basin formation. Solid lines show confined age ranges of the formation of marginal basins by seafloor spreading, and broken lines show rough estimates of the age range. Rhombic marks show the timing of entrapment of the basins (Tamaki and Honza, 1991).

middle Miocene age in the circum-Pacific region is at least 20-30 times that of Green tuff activity near the Japanese Islands at that time, if one considers the total area of Green tuff regions in the circum-Pacific region (Yano, 1985), and total volume of oceanic production rate at back-arc basins during that time (Kaiho and Saito, 1994). Therefore, the total eruption rate at back-arc basins is (8-12) x 1014 g/year which is consistent with that by Kaiho and Saito (1994) (7.4 x 1014 g/year). Therefore, it is estimated that the quantity of Mg removal by the reaction of seawater with volcanic rocks at back-arc basins in the circum-Pacific region during 25-15 Ma is 2.6 i 1 x 1013 g/year (Table 3.1). Annual gain for Ca, K, and Si in seawater due to seawater cycling through back-are basins in Green tuff region in the circum-Pacific region during the early to middle Miocene and those through present-day midoceanic ridges are shown in Table 3.1.

Chapter 3

412 i .............

,j

Waming event Rapid cooling event

CRETACEOUS I PALE~ 70

I

60

EOCENE 50

30

I Pulo

MIOCENE

IOLIGOCENEI 40

20

10 !

Kuril Japan

I t I 1 I I t I

I

,

I I I

'II

! !

l I I I I I I I I I I I

Okinawa Marlana

o n

Pareoe Vela

o

W. Philippine

uJO3

Dalto South China Sulu

nz0

! I

a I

I I

1

I,

'',

I I !

I I

! !

!

I

I

,

Caroline

BB

Manus

BB

I !

' I !

I ! , !

! I !

I 1 I

I

! i I I

I I I I

I I I ! I I I I ;

I I I I I I I 1 ;

I

~

'

-

I I

:

,

I I I I

I

u. m

Celebes Molucca Makassar Andaman

1

~

' //

II

I

i

,

Shikoku

I

g.. ! I

BASIN

Komandosky

I ! !

I

AGE 0 Ma

I i . ~ , . . ~ l

I I ! I I I I I I

Soloman Woodlark Coral Loyalty

<.)

North Fill

p.

Lau-Harve

-iI.-

New

O O3

E

O3 UJ

II

Caladonia

South Fiji Emerald Central Scotia East Scotia Cayman

I

106 km3/m.y.

= E .-U) ~

3

ja. 1

~.~._

o Fig. 3.5. Variations in rates of production of back-arc basin crust. Thickness of each bar shows crust production rate; the scale is the same for all plots. Area of each rectangle shows area of spreading. Warming and cooling events are also shown (Kaiho and Saito, 1994).

Hydrothermal Flux from Back-Arc Basin and Island Arc

413

TABLE 3.1 Annual gain and loss for some constituent in seawater due to seawater cycling through back-arc basins in the Green tuff region in the Circum-Pacific region during the early to middle Miocene, and those through present-day midoceanic ridges (in g/year) (Shikazono, 1994) Green tuff region

Midoceanic ridges

Mg

- 2 . 6 4-1 x 1013

Ca

+4.3 4-1.7 • 1013

K

+4.2 • 1.6 x 1013

Si

+ 1.0 4- 0.4 x 1014

- 2 . 4 • 1013 (Mottl, 1983) - 6 . 5 • 1013 (Wolery and Sleep, 1976) +2.3 x 1013 (Edmond et al., 1979) +94-3 x 1013 (Wolery, 1979) +8.1 • 1012 a +3.9 • 1013 b +0.4 • 1013 (Hart and Staudigel, 1982) +2.3 4-1.2 • 1013 (Holland, 1984) + 1.5 • 1013 (Palmer and Edmond, 1985) +6 4- 8 • 1013 (Wolery and Sleep, 1976)

a Estimation assuming molar exchange of Mg and Ca using 2.4 x 1013 g/year for Mg flux. Estimation assuming molar exchange of 1 mol K and using 2.4 x 1013 g/year for Mg flux.

b

Mg removal by seawater cycling through midoceanic ridges has been estimated by several workers. They vary widely and are: - 2 . 4 x 1013 g/year (Mottl, 1983), - 6 . 5 x 1013 g/year (Wolery and Sleep, 1976), and - 3 . 9 x 103 mol/year (Elderfield and Schultz, 1996). Kaiho and Saito (1994)estimated 2 0 x 106 km3/m.y, and 2 x 106 km3/m.y. for present-day midoceanic ridge crustal production rate and back-arc basin crustal production rate, respectively. If their estimates are correct, Mg removal to midoceanic ridge basalt during early-middle Miocene age is estimated to be 2.6 • 1 x 1013 g/year. Although estimates of annual Mg removal by interaction of circulating seawater with midoceanic ridge basalt are uncertain, it seems likely that Mg removal by seawatervolcanic rock interaction at back-arc basins corresponds to that of Mg removal at midoceanic ridge axis.

3.2. Volatile element (CO2, S, As) flux 3.2.1. CO2 flux

Previous studies demonstrated that the C O 2 fluxes by hydrothermal solution and volcanic gas from midoceanic ridges play an important role in the global CO2 cycle and affect the CO2 concentration in the atmosphere (e.g., Javoy, 1988). However, submarine volcanism and hydrothermal activity occur not only at midoceanic ridges but also at island arc and back-arc basins as already noted. The CO2 concentrations of present-day hydrothermal solutions venting from backarc basins and midoceanic ridges are summarized in Table 3.2. The data show that the CO2 concentrations of hydrothermal solution from back-arc basins and midoceanic

414

Chapter 3

ridges range mostly from 30 to 200 mmol/kg. H20 and from 5 to 30 mmol/kg. H20, respectively. The CO2 concentrations of hydrothermal solutions at Guaymas Basin vary widely and some data show high CO2 concentrations. These high CO2 concentrations and low 313C of fluids (-10.5%o) are considered to be caused by the effect of decomposition and dissolution of organic matters and carbonates in the sediments overlying basalt (Simoneit et al., 1984). The CO2 concentrations of hydrothermal solution from back-arc basins can be also estimated from the fluid inclusion data on Kuroko deposits (Table 3.2). In order to estimate hydrothermal CO2 flux, fluxes of hydrothermal solution into the ocean have to be estimated. As discussed in sections 1.5.3, 2.3 and 2.4.1, the hydrothermal solutions both from back-arc basins and midoceanic ridges are dominantly of seawater origin. Therefore, the fluxes of hydrothermal solution are estimated from seawater cycling rate. This rate is considered to be equal to oceanic production rate times seawater/rock ratio. Kaiho and Saito (1994) estimated the crustal production rate at back-arc basins (Okinawa, Mariana, Andaman, Manus, Woodlark, North Fiji, Lau-Havre, East Scotia and Cayman) based on the spreading rate, thickness of crust and length of ridge axis. Their estimated oceanic crustal production rate is 8.5 x 106 km3/m.y, which is roughly equal to 2.5 x 1015 g/m.y. The seawater/rock ratio can be estimated from the chemical and isotopic compositions of rocks altered by seawater-rock interaction and fresh rocks (Table 3.3). Particularly Mg concentration of altered rocks is useful for the estimation because Mg in seawater removes almost completely into the rocks through the reaction of seawater with rocks at elevated temperatures. Shikazono (1994) summarized the chemical compositions of altered igneous rocks in the Green tuff region of Miocene age in Japan and estimated seawater/rock ratios (by weight) to be 12-18. Oxygen isotopic compositions of altered rocks in the Kuroko mine area can be used to estimate seawater/rock ratio (by weight) to be more than 1 to less than 40 (Green et al., 1983; Shikazono et al., 1995). Therefore, seawater/rock ratio for the discharge zone is assumed to be 2. The average seawater/rock ratio for the submarine hydrothermal system is higher, probably 5-20. Using 2.5 x 1015 g/m.y, as oceanic production rate and 5-20 as seawater/rock ratio and assuming that 30% of oceanic crust interacts with circulating seawater, and the crustal production rate is (0.8-1.1) • 1019 kg/m.y., then the rate of seawater cycling through back-arc basin is estimated to be ( 4 - 2 2 ) x 1019 kg/m.y. Using this value and the CO2 concentration of hydrothermal solution ((0.05-0.3) mol/kg. H20) (Table 3.2), hydrothermal CO2 flux into the ocean is estimated to be (0.2-6) x 1019 kg/m.y. The CO2 flux by hydrothermal solution from midoceanic ridges can be estimated based on the similar procedure mentioned above. The crustal production rate at midoceanic ridges (discharge zone) is 5 x 1022 g/m.y, by Kaiho and Saito (1994) and 4.5 x 1022 g/m.y, by Holland et al. (1996). Seawater/rock ratios for midoceanic ridge hydrothermal systems previously estimated vary widely: Humphris and Thompson (1978), 2-17 (by weight); Wolery and Sleep (1976), 3.5 (by weight); Holland (1978), 10 (by weight). If we accept 5-20 for the seawater/rock ratio, we can calculate the seawater cycling rate at midoceanic ridges as (0.8-4.6)x 1017 g/year. This value is consistent

Hydrothermal Flux from Back-Arc Basin and Island Arc

415

TABLE 3.2 CO2 concentration of hydrothermal solutions venting from midoceanic ridges (MOR) and back-arc basins (BAB) Locality

CO2 (mmolal)

Temperature (~

209 160-200 64-96 34-42 43.4 42.1 11.1-14.4 196 134

320 220 267-278 296-311 238-287 220 285-291 250-280 250-280

5.7-8.0 11-18 2.6-6.5 16-24 more than 9 8-22 3.7-4.5 179-285 4-12 (90-115) 50

273-355 354-381 less than 403 270-315 220 345--400 140-332 28-299 265-276 283-321 328

10 140 350 240 210 230 150 90 160 60 100 130 310 222 835 1001 134 302 223 190 257

300-400

BAB (back-arc basins)

Okinawa Trough JADE CLA South Ens lzu-Bonin Suiyo SM Mid-Mariana Trough Alic Springs South Mariana Trough Forest vent North Fiji Basin Myojinsyo

MOR (midoeeanic ridges) East Pacific Rise 21~ 13~ 9-10~ Guaymas Basin Escanaba Trough Juan de Fuca Endeavour Juan de Fuca Rise S. Cleft. Seq. Juan de Fuca Rise ASHES Juan de Fuca Rise Middle Val Dead Dog Bent Bent Ril Juan de Fucca ASHES (Cl-normal) Juan de Fuca Rise ASHES (Cl-enriched)

Kuroko deposits Uwamuki No. 2 Uchinotai W Fukazawa

Matsumine

Shakanai Kosaka Matsumine Nurukawa

300-400 200-300 200-300 300-400 200-500 200-500 200-300 300-500 200-300 230-270 230-270 230-270 230-270 230-280 230-280 230-280 230-280 230-280

416

Chapter 3

TABLE 3.3 Water/rock ratio in midoceanic ridge and Kuroko hydrothermal system (Shikazono, 1988) 1. Geochemical estimate

Midoceanic ridge 2

3-7 5 7-16 1

1-5 1-5 0.5

87Sr/86Sr 87Sr/86Sr Chemical composition of hydrothermal solution Chemical composition of hydrothermal solution 8180 87Sr/86Sr 818O 818O

Cyprus

5.6 0.75 0.43

87Sr/86Sr 8180 S content

Ku roko 1

>2 >1.4 >2.3

Chemical composition of ore fluids and rocks Mg content Mg content Ca content

2. Geophysical estimate

Midoceanic ridge

3.5 0.5 <6.6 2 2-10 2-2.5

with 3 4-1 x 1016 g/year (axial flux) by Elderfield and Schultz (1996). The seawater cycling rate at the hydrothermal system (axis, flank and off axis) is higher than this value: Wolery and Sleep (1976), 2.9 x 1017 g/year; Holland (1978), 1.0 x 1017 g/year; Holland (1984), 0.5 x 1017 g/year; Lister (1973), (0.6-3.6) x 1017 g/year; Palmer and Edmond (1989), 1.2 x 1017 g/year; Kadko et al. (1995), 2.9 x 1017 g/year; Holland et al. (1996), 4.1 x 1017 g/year. The CO2 concentration of the hydrothermal solution of midoceanic ridges is ca. (0.005-0.01) mol/kg. H20 (Table 3.2). Thus, hydrothermal CO2 flux from midoceanic ridges is estimated to be (2.2-18.0)x 102~215 0 . 0 0 5 - - ( 0 . 4 - 9 ) x 1018 mol/m.y. This is consistent with the estimate of Elderfield and Schultz (1996). The value of (0.4-9.0) x 10 Is mol/m.y, can be compared with the CO2 flux by volcanic gas from midoceanic ridges: Marty and Jambon (1987), 2.2 x 1018 mol/m.y.; Des Marais (1985), 1-8 x 1018 mol/m.y.; Tajika and Matsui (1990), 4.0 x 1018 mol/m.y.; Tajika (1998), 2.0 x 1018 mol/m.y. Javoy et al.'s (1982) estimate (20.3 x 1018 mol/m.y.) is quite different and higher than those by other researchers.

Hydrothermal Flux from Back-Arc Basin and Island Arc

417

TABLE 3.4 Estimates of present-day global fluxes (mol/year) of CO2 to and from the atmosphere (Seward and Kerrich, 1996; Shikazono and Kashiwagi, 1999) Geologic process

Global flux

Chemical weathering Midoceanic ridge system Subaerial volcanism Subaerial + submarine volcanism Anthropogenic Back-arc Subaerial hydrothermal

-6.7 x 1012 +0.7 to 1.5 x 1012 +1.2 to 1.8 x 1012 -+-2to 4 x 1012 +5 x 1014 +5 to 16 x 1012 > 1012

The above argument indicates that hydrothermal C02 flux from back-arc basins is similar to or greater than that from midoceanic ridges, and thus the hydrothermal flux from back-arc basins as well as hydrothermal flux from midoceanic ridges have to be taken into account when we calculate global geochemical CO2 flux. At the convergent plate boundaries, CO2 degasses not only from back-arc basins by hydrothermal solutions but also from terrestrial subduction zones by volcanic gases and hydrothermal solutions. However, the studies on CO2 degassing from terrestrial subduction zones are not many. Seward and Kerrich (1996) have shown that hydrothermal CO2 flux from terrestrial geothermal system (such as Taupo volcanic zone in New Zealand) exceeds 1012 mol/year which is comparable to that of midoceanic ridges (Table 3.4). Sano and Williams (1996) calculated present-day volcanic carbon flux from subduction zones to be 3.1 x 1012 tool/year based on He and C isotopes and CO2/3He ratios of volcanic gases and fumaroles in circum-Pacific volcanic regions. Williams et al. (1992) and Brantley and Koepenich (1995) reported that the global CO2 flux by subaerial volcanoes is (0.5-2.0) x 10 is mol/m.y, and (2-3) x 10 is mol/m.y. (maximum value), respectively. Le Guern (1982) has compiled several measurements from terrestrial individual volcanoes to derive a CO2 flux of ca. 2 x 10 is mol/m.y. Le Cloarec and Marty (1991) and Marty and Jambon (1987) estimated a volcanic gas carbon flux of 3.3 x 1017 mol/m.y, based on C/S ratio of volcanic gas and sulfur flux. Gerlach (1991) estimated about 1.8 x 10 is mol/m.y, based on an extrapolation of measured flux. Thus, from previous estimates it is considered that the volcanic gas carbon flux from subduction zones is similar to or lower than that of hydrothermal solution from back-arc basins.

3.2.2. Causes for high C02 concentration and origin of C02 of hydrothermal solution from back-arc basins The main alteration minerals surrounding Kuroko ore body are K-mica, K-feldspar, kaolinite, albite, chlorite, quartz, gypsum, anhydrite, and carbonates (dolomite, calcite, magnesite-siderite solid solution), hematite, pyrite and magnetite. Epidote is rarely found in the altered basalt (Shikazono et al., 1995). It contains higher amounts of ferrous iron (Fe203 content) than that from midoceanic ridges (Shikazono, 1984).

Chapter 3

418

1 o o ---

--1

-2 -3

-4 I

I

I

150

200

250

I

I

300 350 Temperature('C)

Fig. 3.6. log fco2-temperature diagram showing the univariant equilibrium curves for some gangue minerals. A: 2Ca2A13Si3OIz(OH) (clinozoisite) + 3SIO2 (quartz)+ 2CACO3 (calcite) + 2 H 2 0 - - 3CazA12Si30]0(OH)2 (prehnite) + 2CO2 ( X p i s - - 0 . 3 ) . B: Ca6Si6OI7(OH)2 (xonotlite) + 6CO2 -- 6CaCO3(calcite) + 6SiO2(quartz) + H20. C: CaCO3 (calcite) + TiO2 (rutile) + SiO2 (quartz) --CaTiSiO5 (sphene) + CO2. D: MnSiO3 (rhodonite) + CO2 -- MnCO3 (rhodochrosite) + SiO2 (quartz). E: 3 KAI3Si3010(OH)2 (Kmica) + 4CACO3 (calcite) + SiO2 (quartz) = 2CazA13S3OIz(OH)(clinozoisite) + 3KAISi308 (K-feldspar) + 4CO2 + 2 H 2 0 ( X p i s = 0 . 3 ) . F: 3KAI3Si3OI0(OH)2 ( K - m i c a ) + 4CACO3 (calcite)+ SiO2 (quartz)= 2CazA13Si3OI2(OH) (clinozoisite) + 3 KA1Si308 (K-feldspar) + 4CO2 + 2 H 2 0 ( X p i s = 0 . 2 5 ) . G" 3 FeCO3 (siderite) + (1/2)O2 = Fe304 (magnetite) + 3CO2, C (graphite) + O2 = CO2. H: 3 CaMg(CO3)2 (dolomite) + KAISi308 (K-feldspar) + H 2 0 - - 3CACO3 (calcite) + 3CO2 + KMg3(AISi3OI0)(OH)2 (phlogopite). I: C (graphite) + 02 = CO2, FeS (pyrrhotite) + 1/2S2 = FeS2 (pyrite), 2H2S(aq) -Jr- O2 -- $2 + 2H20(1). J: FeCO3 (siderite) + Fe203 (hematite) = Fe304 (magnetite) + CO2. K: CaA12Si4OIz2H20 (wairakite) + KAISi308 (K-feldspar) + CO2 = CaCO3 (calcite) + KAI3Si3Oj0(OH)2 (K-mica) + 4SIO2 (quartz). L: CaAlzSi4OIz2H20 (wairakite) + CO2 = CaCO3 (calcite) + AIzSi2Os(OH)4 (kaolinite) + 2SIO2 (quartz). M: Ca(A12Si4Ol2.4 H20 (laumontite) + CO2 = CaCO3 (calcite) + AlzSi2Os(OH)4 (kaolinite) + 2 SiO2 (quartz) + 2 H20. Solid star: midoceanic ridge, solid circle: back-arc basin (modified after Shikazono, 1985).

Hydrothermal alteration minerals from midoceanic basalt are analcite, stilbite, heulandite, natrolite-mesolite-scolecite series, chlorite and smectite for zeolite facies, prehnite, chlorite, calcite and epidote for prehnite-pumpellyite facies, albite, actinolite, chlorite, epidote, quartz, sphene, hornblende, tremolite, talc, magnetite, and nontronite for green schist facies, hornblende, plagioclase, actinolite, leucoxene, quartz, chlorite, apatite, biotite, epidote, magnetite and sphene for amphibolite facies (Humphris and Thompson, 1978). The fco2-temperature relationships for the above-mentioned mineral-fluid equilibria are shown in Fig. 3.6. Based on the thermochemical calculations, minerals summarized above and temperatures estimated, we could estimate typical fco2-temperature ranges for hydrothermal solutions from midoceanic ridges and back-arc basins. The analytical data on CO2 in hydrothermal solutions and fluid inclusions and measured temperatures (Table 3.2) are consistent with the thermochemical calculations mentioned above.

Hydrothermal Flux from Back-Arc Basin and Island Arc

419

Ishibashi and Urabe (1995) considered that the high volatile concentrations (C02, etc.) of hydrothermal solutions venting from back-arc basins are due to higher contribution of magmatic fluids containing high concentrations of volatiles than hydrothermal solutions at midoceanic ridges. However, the high concentration of CO2 of hydrothermal solution from back-arc basins does not imply larger contribution of magmatic fluids to hydrothermal solutions at back-arc basins, but is seems more likely from the abovementioned reasons that the alteration minerals buffering fluid chemistry are different for midoceanic ridges and back-arc basins and CO2 in hydrothermal solution was derived from carbonates in altered rocks and not directly from magma. 813C data on the carbonates in altered volcanic rocks in back-arc basin (Kuroko mine area) are -5%~, indicating that most carbon in carbonates were of igneous origin (Shikazono et al., 1995). Stable isotopic studies of 8180 and 8D of hydrothermal solutions venting from back-arc basins show no evidence of contribution of magmatic fluids to the hydrothermal solutions at back-arc basins and midoceanic ridges. As noted already, the stable isotopic data (834S, 813C, 8180, 8D) all indicate that hydrothermal solutions in submarine hydrothermal system in back-arc basins and midoceanic ridges were generated by seawater-rock interaction at hydrothermal conditions. The CO2 concentrations of present-day geothermal waters in terrestrial environment have been also interpreted in terms of the interaction of hydrothermal solutions with country rocks (Giggenbach, 1981; Shikazono, 1978,1985). For example, as noted in section 2.4.3, Shikazono (1985) estimated fco2 for epithermal Au-Ag and base-metal veintype deposits in Japan which formed in terrestrial environments at Miocene-Pliocene age and showed that fco2 is controlled by the alteration minerals (Fig. 3.6). Estimated fcoztemperature range for epithermal Cu-Pb-Zn vein-type deposits are clearly similar to those for the Kuroko and back-arc deposits in which base metals (Cu, Pb, Zn) are concentrated. It is likely that the minerals controlling fco2 of hydrothermal solution at back-arc basins are dolomite, siderite, calcite, hematite, magnetite, graphite, K-mica and kaolinite. Most of these minerals are not found in altered ridge basalt. Among the minerals mentioned above calcite and kaolinite may be important for controlling fco2 of terrestrial geothermal waters. It was cited by Giggenbach (1981) that fco2 (or Xco2, mole fraction of CO2) of terrestrial geothermal waters is controlled by "plagioclase" + CO2 -- calcite + "kaolinite". Berndt et al. (1989) have indicated that acaz+/a2+ and aNa+/aH+ of midoceanic ridge hydrothermal fluids is controlled by clinozoisite, Ca-feldspar, and Na-feldspar. In addition to these assemblages, calcite is in equilibrium with fluids. Therefore, we can derive the fcoz-temperature relationship from the following equilibrium relations. Electroneutrality relation is approximated by, mNa+ -- me1-

(3-2)

The mc1- of midoceanic ridge hydrothermal solution is generally close to that of seawater. If the above argument is correct and the fco2 of hydrothermal solutions from backarc basins is in equilibrium with alteration mineral assemblage including dolomite and calcite, mgg2+/inCa2+ of fluids can be estimated to be 0.03-0.055 from dolomite-calcite-

Hydrothermal Flux from Back-Arc Basin and Island Arc

421

values, we obtain hydrothermal S flux as (0.8-3.2) x 1017 mol/m.y. This estimated value is lower than that from midoceanic ridges and seems to be comparable to volcanic gas flux from terrestrial island arc ((1.5-5) x 1017 mol/m.y.) (Wolery and Sleep, 1976). Sulfur in the sediments and oceanic crust which is derived from seawater subducts to deeper parts. This subduction flux is estimated to be ca. 4 x 1017 mol/m.y. (Shikazono, 1997). Therefore, degassing S flux from back-arc and island arc ((2.3-8.2)x 1017 mol/m.y.) seems to be not different from the subduction flux, although uncertainty of estimated degassing and subduction flux is large. We speculate from the above argument that primordial sulfur degasses from midoceanic ridges even at present time as well as He, because subduction flux to mantle seems to be small. However, we need more detailed study on long-term S cycle including hydrothermal S flux to evaluate this speculation. 3.2.4. As flux

The As (arsenic) concentration of seawater is controlled by input of rivers, sedimentation on the seafloor, weathering of the seafloor, exchange between atmosphere and seawater, volcanic gas input, and hydrothermal input. Previous studies on the geochemical cycle of As have not taken into account the hydrothermal flux of As. Therefore, hydrothermal flux of As from back-arc, island arc and midoceanic ridges to ocean is considered below. As concentrations of submarine and subaerial hydrothermal solutions are summarized in Table 3.5, which clearly shows that the As concentration of hydrothermal solutions from back-arc basins and from subaerial island arc are higher than those from midoceanic ridges. Many analytical data on As concentration of hot springs in subaerial island arc are available (Ellis and Mahon, 1977; Weissberg et al., 1981). These data clearly indicate high concentration of As in these hot springs. An experimental study at 350~ on the interaction between NaC1 solution and graywacke which occurs widely in island arc geologic setting indicates that the final solution contains (0.6-0.7) ppm As (Bischoff et al., 1981). Analytical data on As concentration of hydrothermal solution at back-arc basins are few. Arsenic concentration of hydrothermal solution at Lau Basin is 6.0-8.2 ppm (Foquet et al., 1991). We can also estimate As concentration of hydrothermal solution based on the solubility data on orpiment and realgar because these As-bearing minerals are common in back-arc basin deposits (e.g., Okinawa Trough, Kuroko deposits). Based on the above data and argument, it is inferred that hydrothermal solution venting from back-arc basins contains appreciable amounts of As (1-5 ppm). Using 2.5 x 1019 kg/m.y, as the ocean crustal production rate at back-arc basins (Kaiho and Saito, 1994), 1-10 as seawater/rock ratio, and 1-5 ppm As concentration, we can estimate hydrothermal As flux from back-arc basins to be (1.3-0.13)x 1018 g As/m.y. It seems unlikely that all of the oceanic crust produced interacts with seawater (Holland, 1978). Accepting that 30% of oceanic crust interacts with circulating seawater, hydrothermal As flux is estimated to be (3.8-0.1) x 1011 g As/year. This flux, although

Chapter 3

422 TABLE 3.5

Arsenic concentrations of hydrothermal solutions issuing at midoceanic ridges, back-arc basins and island arcs (Shikazono, 1993) As concentration (ppm) Midocean ridges NGS 21 ~ OBS North SW HG Guaymas

0.002 0.02 0.02 0.03 0.02-0.08

Back-arc basins Lau Basin

6.0-8.2

Island ares Apapel Springs, Kamchatka Broadlands, New Zealand drill 2 Surface Deep aquifer Cheleken, USSR: i Cheleken, USSR: ii Cheleken, USSR: iii Dvukhyurtochnye Springs, Kamchatka Mendeleyev Volcano, Kurile Islands Caldera Springs, Kamchatka Waiotapu, New Zealand, Champagne Pool Wairakei, New Zealand, Hole 44 Pauzhetsk, Kamchatka Steamboat Springs, USA Ngawha, New Zealand

2.5-3.0 8.1 5.5 0.1 0.03 0.5 2.8 2.2 25 4.9 4.8 1.0 2.7 0.2

uncertainty is large, is comparable to riverine flux (7.8 x 10 I~ g/year). Hydrothermal As flux from midoceanic ridges can be also estimated using oceanic crustal production rate (= 5 x 10 22 g/m.y.) (Kaiho and Saito, 1994) and As concentration of hydrothermal solution. This estimated flux is (0.5-0.06) x 10 l~ g/year. This estimated flux is considerably small compared with hydrothermal flux from back-arc basin and riverine flux. The flux of volcanic gas to ocean has not been estimated. Walsh et al. (1979) estimated As flux of volcanic gas to atmosphere to be 2.8 x 109 g/year. Therefore, this flux to ocean is also small. Arsenic removes from basalt to seawater by the weathering of ocean floor basalt. Kawahata and Shikazono (1988) found that the sulfur content of the midoceanic ridge basalt at Galapagos rift decreases from ca. 1,000 ppm to ca. 400 ppm by the seafloor weathering. Average As/S ratio of pyrite is (8.7 4- 3) x 10 -4 (Fleisher, 1955; Utter, 1978; Huerta-Diaz and Morse, 1992). Using this ratio and assuming that all of As in pyrite remove to seawater by the seafloor, weathering and volume of basalt suffered by the seafloor weathering is (4.5-15) x 1014 g/year [this is estimated by assuming that 60-200 m thick basalt is weathered and ocean production rate is 3 x 10 l~ cmZ/year (Deffeyes,

Hydrothermal Flux from Back-Arc Basin and Island Arc

423

TABLE 3.6 Geochemical balance of arsenic in ocean and subduction flux (g/year) (Shikazono, 1993) Input flux to ocean (1) River flux (2) Hydrothermal flux: island arc-back-arc basin axis of midocean ridge (3) Volcanic gas (4) Atmosphere (5) Weathering of ocean floor of basalt

7.8 x 10 l~ (0.2-5.2) x 10 l~ (0.8-1.6) x 1011 2.8 x 109 (max.) 2 . 6 x 109 (2.7-9) x 10 8 Total: (1.0-6.1)x 1011

Output flux from ocean (6) Sedimentation (formation of pyrite) (7) Atmosphere

(1.3-2.9) x 10 ll 1.4 x 10 8 Total: (1.3-2.9) x 1011

Subduction flux

(4.0-8.2) x 10 l~

1970)]. Arsenic flux by the weathering is estimated to be (2.7-9) x 108 g/year. Total As input is a sum of hydrothermal, volcanic gas, riverine, and weathering fluxes which is equal to (1.1-6.1) x 1011 g As/year. Arsenic removal from seawater to sediments is mainly governed by pyrite formation in the seafloor sediments. Production rate of sedimentary pyrite is 2.5 x 1014 g S/year (Holland, 1978). Therefore, As removal by pyrite from seawater is (1.3-2.9) x 1011 g/year. This is the same order of magnitude as As input to ocean by river which is equal to 0.7 • 1011 g/year. Therefore, it is likely that the steady state is maintained with regard to As concentration in seawater. The geochemical balance of As in ocean and subduction flux of As are summarized in Table 3.6.

3.3. Other elemental flux 3.3.1. Hg flux Hg concentration in hydrothermal solution from back-arc basins and midoceanic ridges has not been determined. Experimental study on graywacke-water interaction suggests that the hydrothermal solution interacted with graywacke contains n x 10 -2 ppm Hg (Bischoff et al., 1981). Cinnabar and metacinnabar are not common but were reported from several Kuroko deposits (Urabe, 1974). From the solubility data on cinnabar and metacinnabar (Barnes and Czamanske, 1967), we can place a limit on the Hg concentration of ore fluids to be n • 10 -2 ppm. Using n x 10 -2 ppm concentration and seawater cycling rate at back-arc basins, hydrothermal Hg flux from back-arc

424

Chapter 3

basins is estimated to be n x (108-109 g)/year. Riverine Hg flux is estimated from average Hg concentration of river ((4.0-7.4)x 10 -8 mol/kg. H20) and riverine flux (-- 4.6 x 1022 g/year). This riverine Hg flux is (1.8-3.4) x 109 g/year. Therefore, it is likely that hydrothermal Hg flux from back-arc basins may be important for controlling Hg concentration of seawater, although we need more detailed investigation on the hydrothermal Hg flux.

3.3.2. Mn flux The Mn concentration of hydrothermal solution from back-arc basins varies widely from 12 I~mol/kg.H20 to 7100 I~mol/kg.H20 (Lau Basin, North Fiji Basin) (Gamo, 1995). But, it ranges mostly from 10 txmol/kg. H20 to 300 ~tmol/kg. H20. Using this range and seawater cycling rate (= (0.08-0.8)x 1017 g/year), we obtain Mn flux as (0.08-2.4) x 10 l~ mol/year. Present-day ocean production (back-arc)/ocean production (midoceanic ridge) ratio is about 0.1 (Kaiho and Saito, 1994). According to Elderfield and Schultz (1996), the best estimate of axial flux at midoceanic ridge is (3 4- 1.5) x 1013 kg. HzO/year. If we use this value and the production rate, H20 flux from back-arc basins is estimated at (3 4- 1.5) x 1015 kg-HzO/year. Thus, we estimate hydrothermal Mn flux from back-arc basins to be (1.35-0.15) x 109 mol/year. This hydrothermal Mn flux from back-arc basins is less than the value of riverine Mn flux (0.49 x 1010 mol/year). Hydrothermal Mn flux from midoceanic ridge is estimated as (1.1-3.4) x 10 I~ mol/year (Elderfield and Schultz, 1996) which is one to two order of magnitudes greater than the hydrothermal Mn flux from back-arc basins.

3.3.3. Ba flux Ba concentration of hydrothermal solution from back-arc basins ranges from 5.3 gmol/kg. H20 (North Fiji Basin) to 100 gmol/kg. H20 (Izu-Bonin Suiyo SM) (Gamo, 1995). Assuming that Ba concentration is (20-60) Ixmol/kg. H20 and seawater cycling rate is 1.8 x 1016 g/year, we obtain Ba flux as ( 3 - 6 6 ) x 10 l~ mol/year. This is greater than or comparable to that of midoceanic ridge flux (2.4-13 x 10 l~ mol/year) (Elderfield and Schultz, 1996) and is comparable to or greater than that of riverine Ba flux (1 x 101~ mol/year) (Elderfield and Schultz, 1996).

3.4. Comparison of back-arc hydrothermal flux with midoceanic ridge hydrothermal flux Elderfield and Schultz (1996) estimated midoceanic ridge hydrothermal fluxes using heat and water fluxes estimated by various data (3He/heat, Mg concentration, Sr isotopes, Li isotopes, Ge/Si ratio). Their estimated fluxes are presented in Fig. 3.7 and Table 3.7. Hydrothermal flux from back-arc basins estimated based on the H20 flux which was estimated from oceanic crust production rate and seawater/rock ratio at back-arc

Hydrothermal Flux from Back-Arc Basin and Island Arc

I

425

" 102:1

S0 4 M(

1012 9 Fe Mn LI

l

1"1 ~ ' " g l

Si I 9 " "" K +9 I

, j J

cD

9"

1010 9 jj

=

[ Alk

):102

C~a

I

I

9

108 Co ,'" Ag~ ." "

P

* 106 104 L" " 10 6

l

Se"l

I

10 8

! 10 TM

I

1012

1014

river flux (mol/year) Fig. 3.7. Comparison of hydrothermal fluxes and river fluxes (data from Table 3.7) (Elderfield and Schultz, 1996).

basins and the concentrations of elements in hydrothermal solution from back-arc basins is given above. The following points are inferred from the comparison of back-arc hydrothermal flux mentioned above with midoceanic ridge hydrothermal flux: (1) Back-arc basin hydrothermal flux of most elements is small, compared with midoceanic ridge hydrothermal flux. (2) Back-arc basin hydrothermal fluxes of CO2, As, and Ba are probably higher than the midoceanic ridge hydrothermal fluxes, although we need more detailed investigation. (3) The average chemical compositions of Kuroko ores and those of back-arc deposits suggest that Hg, As, Sb, T1 and Ba are concentrated to the ore fluids responsible for the Kuroko and back-arc deposits, suggesting that these fluxes from back-arc basins are high compared with midoceanic ridge fluxes. (4) Au, Sb and Hg are more enriched into the ores of back-arc basins compared with midoceanic ridge and thus it is likely that back-arc basin hydrothermal flux for these elements is higher than midoceanic ridge hydrothermal flux. However, the concentrations of these elements in back-arc basin hydrothermal solution have not been analyzed. Thus, we need to accumulate analytical data on the concentration of these elements in back-arc basin hydrothermal solution. Further, H20 flux from back-arc basin has to be estimated based on various methods (3He/He, Mg concentration, Li isotope, Sr isotope, Ge/Si ratio) which was argued for midoceanic ridge hydrothermal system by Elderfield and Schultz (1996), but not for back-arc basin by the previous workers.

Chapter 3

426 TABLE 3.7

Comparison of primary axial high-temperature hydrothermal chemical fluxes and fiver chemical fluxes (Elderfield and Schultz, 1996) Element

Li K Rb Cs Be Mg Ca Sr Ba SO4 Alk Si P

Chydrothermafluid (mol/kg) a

Cseawater (mol/kg) a

411-1322 Ix 17-32.9 m 10-33 Ix 100-202 n 10-38.5 n 0 10.5-55 m

26 Ix 9.8 m 1.3 Ix 2.0 n 0 53 m 10.2 m

A1 Mn Fe Co Cu Zn Ag Pb As Se CO2

87 Ix > 8 to >42.6 Ix 0-0.6 m - 0 . 1 to - 1 . 0 m 14.3-22.0 m 0.5 Ix 451-565 Ix 4 - 2 0 Ix 360-1140 IX 750-6470 Ix 22-227 n 9.7-44 Ix 4 0 - 1 0 6 Ix 26-38 n 9-359 n 30-452 n 1-72 n 5.7-16.7 m

CH4 H2 H2S

25-100 Ix 0.05-1 m 2.9-12.2 m

B

87 Ix 0.14 Ix 28 m 2.3 m 0.05 m 2 Ix 416 Ix 0.02 tx 0 0 0.03 n 0.007 Ix 0.01 Ix 0.02 n 0.01 n 27 n 2.5 n 2.3 m 0 Ix 0 m 0 m

Fhydrothermal (mol/year) 1.2 2.3 2.6 2.9 3.0

to 3.9 • 10 l~ to 6.9 • 10 l~ to 9.5 x 104 to 6.0 x 106 to 12 • 105 - 1 . 6 • 1012 9.0 to 1300 • 109

0 >2.4 to 13 x 108 - 8 . 4 x 1011 - 7 . 2 to 9.9 x 101~ 4.3 to 6.6 • 10 II - 4 . 5 x l07 i. 1 to 4.5 • 109 1.2 to 6.0 • 108 1.1 to 3.4 x 10 l~ 2.3 to 19 • 10 l~ 6.6 to 68 x 105 3.0 to 13 • 108 1.2 to 3.2 x 109 7.8 to 11 x 105 2.7 to 110 x 105 0.9 to 140 • l05 3.0 to 220 x 104 1.0 to 12 • l0 II

Frivers (101~ mol/year) 1.4 190 0.037 0.00048 0.0037 530 1200 2.2 1.0 370 3000 640 3.3 5.4 6.0 0.49 2.3 0.011 0.50 1.4 0.0088 0.015 0.072 0.0079

0.67 to 2.4 • 1() l~ 0.3 to !.5 • 1() j~ 0.85 to 9.6 • 1() II

a m = 10-3; tx = 10-6; n = 10 -9.

References Barnes, H.L. and Czamanske, G.K. (1967) Solubilities and transport of ore minerals. In: Barnes, H.L. (ed.), Geochemistry of Hvdrothermal Ore Deposits. New York: Holt, Rinehart and Winston, pp. 334-381. Berndt, M.E., Seyfried, W.E. Jr. and Janeckey, D.R. (1989) Plagioclase and epidote buffering of cation ratios in midocean ridge hydrothermal fluids: Experimental results in and near the supercritical region. Geochim. Cosmochim. Acta, 53, 2283-2300. Bischoff, J.L. and Dickson, F.W. (1975) Seawater-basalt interaction at 200~ 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.J. (1981) Hydrothermal alteration of graywacke by brine and seawater; role of alteration and chloride complexing on metal solubilization at 200~ and 350~ Econ. Geol., 76, 659-676.

Hydrothermal Flux from Back-Arc Basin and Island Arc

427

Brantley, S.L. and Koepenich, K.W. (1995) Measured carbon dioxide emissions from Oldoinyo Lengai and the skewed distribution of passive volcanic fluxes. Geology, 23, 933-936. Deffeyes, K.S. (1970) The axial valley; a steady-state feature of the terrain. Chapter 9 in: Johnson, H. and Smith, B.L. (eds.), The Megatectonics of continents and oceans. New Brunswick, New Jersey: Rutgers U. Press. Delaney, M. and Boyle, E.A. (1986) Lithium in foraminiferal shells: Implications for high-temperature hydrothermal circulation fluxes and oceanic crustal generation rates. Earth Planet. Sci. Lett., 80, 91-105. Des Marais, D.J. (1985) Carbon exchange between the mantle and crust, and its effect upon the atmosphere: Today compared to Archean time. In: Sundquist, E.T. and Broecker, W.S. (eds.), The Carbon Cycle and Atmosphere C02; Natural Variations Archean to Present. Washington, D.C.: Am. Geophysical Union, pp. 602-611. Dick, H.J.B. (1982) The petrology of two back-arc basins of the northern Philippine sea. Am. J. Sci., 282, 644-700. Edmond, J., Measures, C., McDuff, E. et al. (1979) Ridge crest hydrothermal activity and the balances of major and minor elements in the ocean: The Galapagos data. Earth Planet. Sci. Lett., 46, 1-18. Elderfield, H. and Schultz, A. (1996) Midocean ridge hydrothermal fluxes and the chemical composition of the ocean. Annu. Rev. Earth Planet. Sci. Lett., 24, 191-224. Ellis, A.J. and Mahon, W.A.J. (1977) Chemistry and Geothermal Systems. New York: Academic Press. Fleisher, M. (1955) Minor elements in some sulfde minerals. Econ. Geol., g0, 970-1024. Fouquet, Y., von Stackelberg, V., Charlou, J.L., Donval, J.E, Foucher, J.E, Erzinger, J., Herzig, P., Mfihe, R., Wiedicke, M., Soakai, S. and Whitechurch, H. (1991) Hydrothermal back-arc basin; sulfides and water chemistry. Geology, 19, 303-306. Gamo, T. (1995) Wide variation of chemical characteristic of submarine hydrothermal fluids due to secondary modification process after high temperature water-rich interaction: a review. In: Sakai, H. and Nozaki, Y. (eds.), Biogeochemical Processes and Ocean Flux in the Western Pacific. Tokyo: Terra Sci. Publ., pp. 425-451. Gerlach, T.M. (1991) Present-day CO2 emission from volcanoes. EOS (Trans Am. Geophys. Union), 72, 249-251. Giggenbach, W.F. (1981) Geothermal mineral equilibria. Geochim. Cosmochim. Acta, 45, 393-410. Green, G.R., Ohmoto, H., Date, J. and Takahashi, T. (1983) Whole-rock oxygen isotope distribution in the Fukazawa-Kosaka area, Hokuroku district, Japan, and its potential application to mineral exploration. Econ. Geol. Mon., g, 395-411. Hajash, A. and Chandler, G.W. (1981) An experimental investigation of high-temperature interactions between seawater and rhyolite, andesite, basalt and peridotite. Contrib. Mineral. Petrol., 78, 240-254. Hart, S.R. and Staudigel, H. (1982) The controls of alkalines and uranium in seawater by ocean crust alteration. Earth Planet. Sci. Lett., 58, 202-213. Holland, H.D. (1978) Chemistry of Atmosphere and Ocean. New York: John Wiley and Sons. Holland, H.D. (1984) The Chemical Evolution of the Atmosphere and Oceans. Princeton, N.J.: Princeton U. Press. Holland, H.D., Horita, J. and Seyfried, W.E. Jr. (1996) On the seawater variations in the composition of Phanerozoic marine potash evaporites. Geology, 24, 993-996. Holser, W.T. and Kaplan, I.R. (1966) Isotope geochemistry of sedimentary sulfates. Chem. Geol., 1, 93-135. Huerta-Diaz and Morse, J.W. (1992) Pyritization of trace metals in anoxic marine sediments. Geochim. Cosmochim. Acta, 56, 2681-2702. Humphris, S.E. and Thompson, G. (1978) Hydrothermal alteration of oceanic basalts by seawater. Geochim. Cosmochim. Acta, 42, 107-125. Javoy, M. (1988) Carbon geodynamic cycle revised. Chem. Geol., 70, 39. Javoy, M., Pineau, F. and Allegre, C.J. (1982) Carbon geodynamic cycle. Nature, 300, 171-173. Ishibashi, J. and Urabe, T. (1995) Hydrothermal activity related to arc-back arc magmatism in the Western Pacific. In: Taylor, B. (ed.), Back-arc Basins Tectonics and Magmatism. New York: Plenum Publ., pp. 451-496.

428

Chapter 3

Kadko, D., Baross, J. and Alt, J. (1995) The magnitude and global implications of hydrothermal flux. In: Humphris, S.E., Zierenberg, R.A., Mullineauz, L.S. and Thompson, E. (eds.), Seafloor Hydrothermal Systems. Washington, D.C.: Am. Geophysical Union, pp. 446-466. Kaiho, K. and Saito, S. (1994) Oceanic crust production and climate during the last 100 Myr. Terra Nova, 6, 376-384. Kawahata, H. and Shikazono, N. (1988) Sulfur isotope and total sulfur studies of basalts and greenstones from Hole 504B, Costa Rica Rift: Implications for hydrothermal alteration. Can. Mineral., 26, 555-565. Konda, T. (1974) Bimodal volcanism in the northeast Japan arc. Geol. Soc. Japan J., 80, 81-89. Le Cloarec, M.-F. and Marty, B. (1991) Volatile fluxes from volcanoes. Terra Nova, 3, 17-27. Le Guern, E (1982) Les d6bits de C02 et de 802 volcaniques dans l'atmosphbre. Bull. Volcanol., 45(3), 197-202. Lister, C.R.B. (1973) Hydrothermal convection at seafloor spreading centers; source of power or geophysical nightmare? Geol. Soc. Am. Abst. Programs, 5, 74. Marty, B. and Jambon, A. (1987) C/3He in volatile fluxes from the solid Earth: Implications for carbon geodynamics. Earth Planet. Sci. Lett., 83, 16-26. Mottl, M.J. (1983) Metabasalts, axial hot springs and the structure of hydrothermal systems at midoceanic ridges. Geological Soc. Am. Bull., 94, 161-180. Mottl, M.J. and Holland, H.D. (1978) Chemical exchange during hydrothermal alteration of basalt by seawater. I. Experimental results for major and minor components of seawater. Geochim. Cosmochim. Acta, 42, 1103-1115. Ozawa, A. (1963) Neogene orogenesis, igneous activity and mineralization in the central part of northeast Japan. I. On the Neogene igneous activity. J. Japan. Assoc. Mineral. Petrol. Econ. Geol., 50, 167-184 (in Japanese with English abst.). Palmer, M.R. and Edmond, J.M. (1989) The strontium isotope budget of the modern ocean. Earth Planet. Sci. Lett., 92, I 1-21. Reed, M.H. (1983) Seawater-basalt reaction and the origin of greenstones and related ore deposits. Econ. Geol., 78, 466-485. Sano, Y. and Williams, S.N. (1996) Fluxes of mantle and subducted carbon along convergent plate boundaries. Geophys. Res. Lett., 23, 2749-2752. Seyfried, W.E. Jr. (1987) Experimental and theoretical constraints on hydrothermal alteration processes at midocean ridges. Ann. Re~: Earth Planet. Sci., 15, 317-335. Seward, T.M. and Kerrich, D.M. (1996) Hydrothermal emission from the Taupo volcanic zone, New Zealand. Earth Planet. Sci. Lett., 139, 105-113. Shikazono, N. (1978) Possible cation buffering in chloride rich geothermal waters. Chem. Geol., 23, 234-259. Shikazono, N. (1984) Compositional variations in epidote from geothermal areas. Geochem. J., 18, 181-187. Shikazono, N. (1985) Gangue minerals from Neogene vein-type deposits in Japan and an estimate of their CO2 fugacity. Econ. Geol., 80, 754-768. Shikazono, N. (1988) Difference in elements enriched into Kuroko and ridge hydrothermal ore deposits. Ocean Science, 20, 217-222 (in Japanese). Shikazono, N. (1993) Influence of hydrothermal flux on arsenic geochemical balance of seawater. Chikyukagaku (Geochemistry), 27, 135-139 (in Japanese). Shikazono, N. (1994) Hydrothermal alteration of green tuff belt, Japan: Implications tor the influence of seawater/w)lcanic rock interaction on the seawater chemistry at a back arc basin. The Island Arc, 3, 59-65. Shikazono, N. (1997) Geochemistry ~?["Earth System. Tokyo: U. Tokyo Press, 319 pp. (in Japanese). Shikazono, N. and Kashiwagi, H. (1999) Carbon dioxide flux due to hydrothermal venting from back-arc basin and island arc and its influence on global carbon dioxide cycle. 9th Annual V.M. Goldschmidt Conference, August 22-27, Harvard, Abstr., p. 272. Shikazono, N. and Kawahata, H. (1987) Compositional differences in chlorite from hydrothermally altered rocks and hydrothermal ore deposits. Can. Mineral., 25, 465-474. Shikazono, N., Utada, M. and Shimizu, M. (1995) Mineralogical and geochemical characteristics of hydrothermally altered basalt in Kuroko mine area, Japan: Implications for the evolution of back arc basin hydrothermal system. Applied Geochemistry, 10, 621-642.

Hydrothermal Flux from Back-Arc Basin and Island Arc

429

Simoneit, B.R.T., Philip, R.E, Jeden, ED. and Galimov, E.M. (1984) Organic geochemistry of Deep Sea Drilling Project sediments from the Gulf of California - Hydrothermal effects on unconsolidated diatomic ooze. Org. Geochem., 17, 173-205. Sugimura, A. and Uyeda, S. (1973) Island Arc. Iwanamishoten. Sugimura, A., Matsuda, T., Chinzei, K. and Nakamura, K. (1963) Quantitative distribution of late Cenozoic volcanic materials in Japan. Bull. Volcanol., 26, 125-140. Tajika, E. (1998) Climate change during the last 150 million years: reconstruction from a carbon cycle model. Earth Planet. Sci. Lett., 160, 695-707. Tajika, E. and Matsui, T. (1990) Evolution of terrestrial proto-CO2 atmosphere with thermal coupled history of the Earth. Earth Planet. Sci. Lett., 113, 251-266. Tamaki, K. and Honza, E. (1991) Global tectonics and formation of marginal basins: Role of the western Pacific. Episodes, 14, 224-230. Tamura, M. (1982) Alteration minerals and mineralization in the Shakanai Kuroko deposit, Akita Prefecture. Mining Geology, 32, 379-390. Tsunakawa, H. and Takeuchi, A. (1986) Paleo-stress field and igneous activity of Japanese Island. In: Taira, A. and Nakamura, K. (eds.), Formation of the Japanese Island. Iwanamishoten, pp. 201-208 (in Japanese). Urabe, T. (1974) Mineralogical aspects of the Kuroko deposits in Japan and their implications. Mineralium Deposita, 9, 309-324. Utter, T. (1978) Morphology and geochemistry of different pyrite types from the upper Witwatersrand system of the Klerksdorp Goldfield South Africa. Geol. Rundschau, 67(2), 774-804. Von Damm, K.L., Edmond, J.M. Measure, C.I., Walden, B. and Weiss, R.E (1985a) Chemistry of submarine hydrothermal solutions at 21~ East Pacific Rise. Geochim. Cosmochim. Acta, 49, 2197-2220. Von Damm, K.L., Edmond, J.M., Measure, C.I. and Grant, B. (1985b) Chemistry of submarine hydrothermal solutions at Guaymas Basin, Gulf of California. Geochim. Cosmochim. Acta, 49, 2221-2237. Walsh, ER., Duce, R.A. and Fashing, J.C. (1979) Considerations of the enrichment sources, and flux of arsenic in the troposphere. J. Geophys. Res., 84, 1719-1726. Weissberg, B.G., Browne, P.R.L. and Seward, T.M. (1981) Ore metals in active geothermal systems. In: Barnes, H.L. (ed.), Geochemistry of Hydrothermal Ore Deposits. New York: Wiley, pp. 738-780. White, D.E. (1967) Mercury and base-metal deposits with associated thermal and mineral waters. In: Barnes, H.L. (ed.), Geochemistry of Hydrothermal Ore Deposits. New York: Holt, Rinehart and Winston, pp. 575-631. Williams, T.J., Schaefer, S.J., Calvache, U.M.L. and Lope, Z.D. (1992) Global carbon dioxide emission to the atmosphere by volcanoes. Geochim. Cosmochim. Acta, 56, 1765-1770. Wolery, T.J. (1979) Some chemical aspects of hydrothermal processes at midoceanic ridges: A theoretical study. I. Basalt-seawater reaction and chemical cycling between the oceanic crust and oceans. II. Calculation of chemical equilibrium between aqueous solutions and minerals. Ph.D. Thesis, Northwestern U., Evanston, Illinois (unpublished). Wolery, T.J. and Sleep, N.H. (1976) Hydrothermal circulation and geochemical flux at midocean ridges. J. Geol., 84, 249-275. Wood, D.A., Jordon, J.L., Marsh, N.G., Tarney, J. and Greuil, M. (1980) Major and trace element variations drilled in basalts from the North Philippine Sea during Deep Sea Drilling Project Leg 58: A comparative study of back-arc basin basalts with lava series from Japan and midocean ridges. Deep Sea Drilling Project, Initial Reports, 58, 827-894. Yano, T. (1985) Proposal of "Pacific Problems" and its development. Chikyukagaku (Earth Science), 39, 313-318.