The petrology of Tambora volcano, Indonesia: A model for the 1815 eruption

Journal of Volcanology and Geothermal Research, 27 (1986) 1--41 Elsevier Science Publishers B.V., Amsterdam -- Printed in The Netherlands

THE PETROLOGY OF TAMBORA FOR THE 1815 ERUPTION

VOLCANO,

INDONESIA:

1

A MODEL

J. FODEN

Department of Geology, University of Adelaide, G. P.O. Box 498, Adelaide, S.A. 5001, Australia (Received September 20, 1984; revised and accepted July 10, 1985)

ABSTRACT Foden, J., 1986. The petrology of Tambora volcano, Indonesia: a model for the 1815 eruption. J. Volcanol. Geotherm. Res., 27: 1--41. Tambora is an active volcano in the east Sunda Arc and is well known for its catastrophic eruption in April 1815. Its lavas are of unusual, moderately undersaturated, K~O-rich types, ranging from ne-trachybasalt to ne-trachyandesite. They evolved under conditions of high oxygen fugacity about two log units below the M-H buffer, and crystallised leucite at low pressures. The products of the 1815 eruption are black, glassy, biotite-bearing, ne-trachyandesites together with scoria, pumice and tuff of the same composition. The parent trachybasalt liquid probably contained about 3 wt.% H20 and the tra~hyandesites nearly 6 wt.%. The model presented here suggests that the 1815 eruption, which is known to have followed a lengthy period of inactivity, was the result of gradual cooling of a hydrous magma in a closed, high-level magma chamber emplaced at depths between 1.5 and 4.5 kin. Eruption followed the exsolution of a high-pressure aqueous fluid phase formed during cooling and crystallisation of the magma (second boiling). By the time eruption took place, overpressures of about 4--5 kbar had been generated and temperatures in the magma chamber ranged from about 850°C to 700°C. The catastrophic eruption took place after failure of the r o o f and the discharge of the vapour phase may have reached velocities of 650 m s -I. Calculations indicate that the volume of magma involved in the eruption was of the order of 33 km 3 and this may have evolved about 1.2 x 1027 ergs of energy during the eruption, a blast equivalent to 30,000 megatons of TNT.

INTRODUCTION T a m b o r a is a n a c t i v e v o l c a n o o n t h e I n d o n e s i a n i s l a n d o f S u m b a w a a n d i t is r e n o w n e d f o r i t s e x t r a o r d i n a r i l y v i o l e n t e r u p t i o n in 1 8 1 5 ( P e t r o e s c h e v s ky, 1949; Stothers, 1984). Estimates of the size of this eruption have been m a d e b y s e v e r a l w o r k e r s a n d m o s t p l a c e i t as p e r h a p s t h e l a r g e s t v o l c a n i c e r u p t i o n i n h i s t o r i c t i m e (e.g. Y o k o y a m a , 1 9 5 7 ) . T h i s c o n c l u s i o n is a l s o c o r r o b o r a t e d b y t h e r e c e n t w o r k o f S e l f e t al. ( 1 9 8 4 ) . Petroeschevsky (1949) and Stothers (1984) have summarised the early l i t e r a t u r e d e a l i n g w i t h t h e e r u p t i v e e v e n t . A c c o u n t s q u o t e d in P e t r o e s c h e v s -

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ky's paper suggest that the massive eruption which t o o k place between April 10 and April 12, 1815 was preceded b y at least several centuries of dormancy. Petroeschevsky (1949) also described the petrography of samples of lava and pumice he collected on the northwestern slope of the volcano. These he grouped into olivine basalts, leucite basanites and biotite-bearing glassy bombs. More recently, F o d e n and Varne (1980, 1981) have published some geochemical and petrographic data from the suite of rocks collected by the author in 1976. Their study showed that the Tambora suite was an undersaturated, moderately K20-rich association ranging from ne-trachybasalt to ne-trachyandesite. F o d e n and Varne (1980) found that this suite fell geochemically between the very undersaturated, K20-rich lavas from some other Sumbawan volcanoes (Soromundi and G. Sangenges) and the relatively K20-poor calc~lkaline suites of the region (e.g. Rindjani volcano, L o m b o k island; Foden, 1983). This paper has two main aims; the first to document the petrology and geochemistry of the highly unusual y o u n g lavas from this most important volcano and the second to discuss the 1815 eruption and its causes. These two aims are complimentary as the crystallisation history of the 1815 liquids can only be interpreted in the light of knowledge of the pre-1815 Tambora magmas. TECTONIC SETTING The geology and tectonics of Indonesia have been discussed and interpreted by Van Bemmelen {1949) and Hamilton (1979). The geological setting of this part of the east Sunda Arc was also summarised b y F o d e n {1983) and by F o d e n and Varne (1980). More recently, discussions focussing on major problems such as the timing of arc collision with the Australian continental shelf, the contentious tectonic position of Sumba island (Fig. 1) and the possibility of back-arc rifting and subduction reversal (in the Flores Sea to the north of Sumbawa and Flores) form the basis of a number of papers (e.g. Silver et al., 1983; Chamalaun et al., 1984; v o n d e r Borch et al., 1984). Very briefly, S u m b a w a is in the active inner volcanic arc of the Sunda-Banda island-arc system. Tambora volcano lies some 340 km north of the Java Trench system (Hamilton, 1979) and is a b o u t 180--190 km above the upper surface of the active north-dipping subduction zone (Cardwell and Isacks, 1978). The crust in this region is quite thin and velocity profiles suggest it m a y be oceanic in character to the north of Sumbawa {e.g. Silver et al., 1983). In this sector of the arc, in contrast to the situation further east, material underthrust at the Java Trench is of unequivocal oceanic crustal origin. In the immediate vicinity of Sumbawa, with its unusual concentration of very alkaline lavas erupted from several volcanoes (Foden and Varne, 1980) there is evidence of tectonic complexity in the anomalous concentration of shallow and intermediate focus earthquakes particularly south of the inner

arc. The offset southwards of the active volcano line to the east of Sumbawa on Flores island (Fig, 1) has led to the postulation of a major cross-arc fracture zone (Audley-Charles, 1975; Carter et al., 1976). This structure is still hypothetical yet it is certainly true that there is a marked change to tholeiitic lavas on Flores (Wheller et al., 1983) in contrast to more alkalic types from Sumbawa. THE GEOLOGY OF SUMBAWA Aspects of the geology of S u m b a w a have been summarised by Van Bemmelen (1949), Sudradjat et al. (1975), F o d e n and Varne (1980) and Alzwar et al. (1981). The oldest rocks recognised on the island occur in the south and west and are gently folded Miocene sequences of limestone, sandstone and conglomerate with some acid and intermediate volcanic and intrusive rocks. Above these are flat-lying Pliocene to Holocene deposits of limestone, sandstone and conglomerate with intercalated volcanic rocks and large volcanoes. Quaternary volcanic rocks show wide-ranging compositions from highly undersaturated very K20-rich types through to oversaturated c a l c ~ k a l i n e examples (Foden and Varne, 1980; Alzwar et al., 1981). Tambora itself forms the north end of the Sanggar peninsular (Fig. 1) in north-central Sumbawa. It is a b o u t 60 km in diameter at sea level and has a present elevation of 2850 m. Prior to the 1815 eruption its estimated elevation was a b o u t 4000 m (Petroeschevsky, 1949). 2~#o km offshore to the northwest of the main cone is a small adventive cone 3 km in diameter (Satonda island) composed of lava flows of composition similar to some of the lavas of the main part of Tambora itself. Samples described in this study (Tables 1, 2 and 3) were collected from Satonda island, from the NW flanks of Tambora and from the caldera rim. Most of those from the NW flank were in situ and represent flows between 52,000 (K-At) and 5,000 (radio carbon) years old (Barberi et al., 1983). Material from the caldera rim is black, scoriaceous or glassy lava, pumice and tuff derived from the 1815 eruption. This material also includes xenoliths and loose blocks of medium-grained alkali gabbro and syenite. THE PETROGRAPHY

OF THE TAMBORA

LAVAS AND INTRUSIVE ROCKS

The rocks sampled for this study fall into four distinct petrological groups: (1) High MgO ne-trachybasalts (eg PS8, PS2, Tables 1 and 2). These are holocrystalline porphyritic lavas. They are composed of phenocrysts of olivine, clinopyroxene and plagioclase and microphenocrysts of magnetite, in a groundmass of pla~oclase, alkali feldspar (sanidine-anorthoclase), augite, olivine, magnetite, apa~te and leucite. In general the proportion of phenocrysts of olivine and clinopyroxene is greater in the samples from S a t o n d a island as is the relative proportion of olivine with respect to clinopyroxene.

(2) Low magnesian ne-trachybasalts (e.g. T5, T30, Tables 1 and 2). These are petrographically quite similar to the high MgO ne-trachybasalts. They are also holocrystalline porphyritic lavas with phenocrysts of olivine, clinopyroxene (salite -- Ca-rich augite) and plagioclase. They also have microphenocrysts of magnetite, apatite and occasionally leucite and differ from the first group in having a higher proportion of plagioclase phenocrysts relative to those of olivine and clinopyroxene, but have similar groundmass mineralogies. (3) Ne-trachyandesites (banakites or phonolites) (e.g. T32, T43, T27; Tables 1 and 2). These form the product of the 1815 eruption. They are highly unusual rock types in an island-arc environment being relatively silicarich (54--57 wt.% SiO2) yet quite strongly undersaturated. The only well known suite with rocks directly equivalent to these, is that of the Roman Region in southern Italy. Here, Appleton (1972) describes the high-K rocks from the Roccamonfino Volcano and notes that there are (relatively) high and low K20 series. Trachyandesites from Tambora are very like some lavas of his low K series. Most of these ne-trachyandesites have a glassy matrix (Table 11) with phenocrysts of biotite, olivine, clinopyroxene, plagioclase, apatite and magnetite. Microlites of sanidine occur. The plagioclase phenocrysts are largely unzoned and of quite-calcic composition (Ans0_ss) with marked corrosion, numerous glass inclusions and fine sodic rims. Clinopyroxene also contains glassy inclusions. The glassy trachyandesites are entirely free of feldspathoids of any kind, though a few samples were more crystalline and in these there is groundmass leucite. In these rocks the leucites occur as rather large indistinct grains,

TABLE 1 M o d a l c o m p o s i t i o n s o f s o m e T a m b o r a lavas a n d i n t r u s i v e s S a m p l e N o . : PS8 clinopyroxene olivine biotite magnetite plagioclase apatite K-feldspar groundmass glass

11 4.3 0.2 3 24.5 -. 57 .

T17

T13

T5

T30

T32

T4

T6

7.3 2.0 -1.7 14.5 --

10.2 3.5 0.5 4.2 4.4 2.4 . 37.5 .

2.5 0.8 -0.2 38.5 -. 58 .

6 1.2 0.2 0.5 38 0.2

3 -2 1.6 12 0.4

54

-81

10.6 6.5 5.7 3.6 52 1.0 2.04 ---

5.5 -6.6 7.2 39 1.0 40.5 ---

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. 74.5

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P 5 8 , T 1 7 a n d T 1 3 are M g O - r i c h t r a c h y b a s a l t s . T5, T30 M g O - p o o r trachybasalts. T32 1815 trachyandesite. T4, T6 syenitic intrusive rocks.

.

Analyses are formed using are ne-traehy T43 are 1815

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Diwo

ne

an

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0.75 0.63 0.52 0.60 13.12 24.11 24.94 3.76 8.69 6.34 1.53 5.71 1.52 7.68 1.86 1.09

10.52 23.88 26.73 1.96 9.94 7,48 1.46 6.76 1.45 7.48 1.86 0.81

19.11 24.17 23.67 5.70 10.53 7.38 2.26 4.79 1.61 7.25 1.73 1.14

1.37 0.65 0.51 0.42

49.21 17.27 5.00 5.23 0.20 5.71 10.50 4.10 1.71 0.91 0.49

T8

15.78 17.09 25.38 6.79 9.92 7.08 1.97 4.24 1.30 8.03 1.84 0.98

1.02 0.64 0.48 0.76

48.26 17.95 5.54 5.14 0.21 5.27 10.46 3.50 2.67 0.97 0.42

T12

18.32 19.83 22.36 7.32 8.78 5.89 2.24 4.53 1.90 6.48 1.75 0.91

0.97 0.62 0.49 0.79

49.83 18.03 4A7 5.19 0.20 4.96 9.26 3.94 3.10 0.92 0.39

T2

17.08 26.54 t9.59 4.69 9.46 6.43 2.29 3.61 1.41 6.67 1.65 0.93

0.59 0.60 0.48 0.68

5!.49 17.15 3.78 5.62 0.18 4.65 9.04 4.16 2.89 0.87 0.40

T13

21.75 22.99 27.32 7.35 4.86 3.51 0.90 3.25 0.92 5.05 1.25 1.09

0.89 0.64 0.48 0.85

51.48 21.10 3.48 3.16 0.14 3.27 8.47 4.32 3.68 0.66 0.47

T5

0.69 0.56 0.45 0.89

53.33 20.26 2.57 3.75 0.15 2,76 6.72 5.00 4.47 0.64 0.52

T10

21.04 26.42 21.30 25.62 25.75 19.64 8.72 9.04 6.33 4.30 3 + 9 0 2.58 2.07 1.50 2.94 3.01 1.72 1.92 4.00 3.73 1.29 1.22 1 + 1 4 1.21

0.66 0.58 0.47 0.81

51.89 20.56 2.76 4.04 0.15 3.25 8.89 4.42 3.56 0.68 0.49

T30

29.73 32.18 18.32 3.47 2.21 1.47 0.58 3.70 1.60 4.41 1.31 1.26

0.80 0.57 0.45 1.10

55.09 19.66 2.51 3.70 0.18 2.71 5.47 4.56 5.03 0.69 0.54

T32

29.02 29.02 17.34 6.60 3.03 1.45 1.54 3.31 3.88 2.46 1.08 1.33

0.74 0.47 0.52 1.01

54.72 19.68 1.70 4.67 0.18 2.48 5.71 4.87 4.91 0.57 0.57

T27

30.02 32.68 18.05 3.66 2.19 1.59 0.40 3.25 0.90 4.96 1.27 1.28

0.79 0.60 0.43 1 +09

55.17 19.78 3.42 2.81 0.18 2.50 5.42 4.66 5.08 0.67 0.55

T23

3].85 33.71 15.71 4.89 1.67 1,10 0.45 3.07 1.40 4.03 1.16 1.14

0.76 0.57 0.42 1.07

56.11 19.90 2.78 2.85 0,18 2.20 4.62 5.05 5.39 0.61 0.49

T43

19.33 19.85 26.24 5.34 7.41 4.77 2.15 4.55 2.26 5.58 1.67 1.12

0.74 0.60 0.48 0.93

50.03 18.93 3.85 "5.10 0.19 4.52 9.50 3.51 3.27 0.88 0.48

T4

21.93 28.25 27.53 3.72 4.05 3.50 -2.16 -1.53 1.33 1.23 5.40

0.70 0.83 0.43 0.89

52.19 20.93 6.45 0.91 0.19 2.64 8.20 4.15 3.71 0.70 0.53

T6

quoted normalised anhydrous. H20+ values are pre-normalisation figures. Major- and trace-element analyses were perstandard X.R.F+ procedures. Rocks P88 to T13 are ne-traehybasalts. T5 to T10 are low Mg ne+trachybasalts. T32 to T43 andesite (phonolites or " b a n a k i t e " ) T4 and T6 are intrusive nodules of alkali gabbro (shonkinite)+ T32, T27, T23 and eruption products.

H,O Mg/(Mg + Fe 2+ ) Mg/(Mg + Fe) K~O/Na20

P,O+

MgO CaO Na,O K,O TiO,

49.31 17.58 4.12 6.03 0.20 5.82 9.84 3.67 2.22 0.98 0.47

PS2

1.20 0.69 0.56 0.55

49.21 17.07 4.55 5.39 0.18 6.88 10.65 3.25 1.78 0.98 0.35

SiO, A1,O3 Fe,O~

FeO MnO

PS8

Sample No.:

Representative X.R.F. analyses of T a m b o r a lavas and intrusive nodules

TABLE 2

312 0.06 26 3.9

23 55 28

43 49 20

389 0.04 29 3.3

59 977 509 92 3.5 23.6 24.2 275

38 896 465 69 2.4 20.7 33.9 307

PS2

384 0.03 21 4.7

37 1326 1105 102 4.8 21.7 25 285 43 31 73 24 150

T8

264 0.07 26 3.7

84 1234 834 88 3.4 23.6 24 338 8 17 55 23 480 32

T12

306 0.09 23 4.1

84 1113 932 97 4.2 23.5 21 311 10 13 65 18 465

T2

279 0.09 25 4

93 1031 791 104 4.3 25 21 255 7 15 64 26

T13

Tamboralavas andintr~iveroc~

297 0.07 20 5.9

103 1441 990 112 5.5 19 12 166 10 13 61 23 70

T5

284 0.07 21 6.0

104 1426 986 112 5.4 18.7 11 159 30 19 65 26 280

T30

285 0.11 21 6.0

130 1234 1171 132 6.3 22.1 8 126 4 7 74 20 475

T10

N o t e : m a j o r - e l e m e n t s analyses given in T a b l e 2. All e l e m e n t s d e t e r m i n e d b y X . R . F . e x c e p t S w h i c h was d e t e r m i n e d b y i g n i t i o n a n d t i t r a t i o n .

K/Rb Rb/Sr Zr/Nb Zr/Y

Rb Sr Ba Zr Nb Y Sc V Cr Ni Ce Nd C1 S

PS8

~ace~lementcon~ntofselected

TABLE3

316 0,13 18 5.6

360

136 1085 1209 145 7.2 26 10 139 4 4 81 29

T32

304 0.12 20 5.6

134 1110 1210 146 7.3 26.2 10 151 5 5 77 24 1025

T27

348 0.11 21 5.9

121 1091 1203 148 7.2 25.2 8 137 4 6 79 27 1200 140

T23

309 0.14 20 6.0

145 1001 1210 157 7.8 26.0 8 112 4 7 77 26 1015 190

T43

271 0.08 23 4.1

100 1283 780 83 3.6 20.2 18 266 13 12 59 25 735

T4

291 0.07 20 5.3

106 1483 941 118 5.8 22.3 11 166 4 9 70 30 660 270

T6

--3

packed with inclusions of the phenocryst and microphenocryst phases and appears to be a quench product. (4) Coarse Grained Rocks (e.g. T6, T4; Tables 1 and 2). These show a range of chemical compositions, very like although not identical to that of the trachybasalt suite. They range from relatively mafic to quite felsic types. Texturally, the intrusive rocks are somewhat porphyritic with phenocryst assemblages the same as those of the basalts (olivine, clinopyroxene, plagioclase, magnetite and apatite), but with interstitial growth of alkali feldspar, sodic plagioclase and biotite. The plagioclase phenocrysts like many of those from the lavas are largely of very calcic composition (Ans0--Anss) but in these there are broad zoned sodic rims (An~s--An20). They have biotite as phenocrysts which appear later in the crystallisation sequence. Biotite of more Fe-rich composition also occurs as reaction rims around olivine and magnetite. In some of these rocks (e.g. T6) interstitial alkali feldspar is intergrown with sodalite in a graphic texture perhaps indicating eutectoid and hence solidus crystallisation conditions. MINERAL COMPOSITIONS

The compositions of minerals tabulated in this paper were determined by the author using the University of Adelaide's JEOL 733 electron microprobe with attached Kevex energy-dispersive system. An

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Or

(1) Feldspar Plagioclase is present as a phenocryst in all rocks from this suite. Relative to the ferromagnesian phases the proportion of plagioclase is least in the most magnesian trachybasalts (e.g. PS8, Tables 1 and 2) and increases through the trachybasalt series. The plagioclase phenocrysts in b o t h trachybasalts and trachyandesites include the same very calcic compositions in the range Ans0--Ang0 (Table 4, Fig. 2). Plagioclase phenocrysts have corroded cores or zones often riddled with glass inclusions. Outer parts of the phenocrysts may be finely zoned, ranging o u t to a b o u t An60 (Table 4, Fig. 2). The trachyandesite phenocrysts are at least as calcic as those of the trachybasalts (Table 4, Fig. 2) and in many cases they are highly corroded with only the very thinnest rims of more sodic plagioclase (deposited on quenching during eruption). These plagioclases include large, glass-filled, re-entrant cavities. Plagioclase and alkali feldspar coexist in the groundmass of the trachybasalts. In the more mafic of these, plagioclase is more calcic and alkali feldspar more sodic than are those from the low MgO trachybasalt groundmass (Fig. 2). In fact it is c o m m o n l y the case that the feldspar of the groundmass of the more mafic trachybasalts is in fact a single phase sanidine-anorthoclase indicative of the relatively higher quenching temperature of these rocks compared with the less magnesian ones. (e.g. Carmichael et al., 1974, p. 220). The intrusive rocks on the other hand have much more sodic plagioclase and calcium-free K-feldspar in their interstices (Fig. 2, Table 4) indicative of quite low temperatures (Carmichael et al., 1974).

(2) Clinopyroxene Clinopyroxene is present in all rocks. Representative analyses are given in Table 5 and their compositions in terms of Ca-Mg-Fe end members are plotted in Fig. 3. Most phenocrysts of the entire range of lavas are Ca-rich augites with moderate amounts of Ti and A1 and low Na contents. Mg/(Mg + Fe) ratios are mostly between 0.7 and 0.75. Groundmass clinopyroxene is more Al-poor than the phenocrysts but more Na- and Ti-rich perhaps indicating the presence of both acmite and perhaps a hypothetical NaTiAISiO6 molecule. A possible reaction favouring the production of such a c o m p o n e n t could be: Fe2TiO4 + NaAISi206 ~ NaTiA1SiO6 + Fe2SiO4 Ti m a g n e t i t e

cpx

cpx

olivine

which is favoured to the right at lower temperatures and pressures and is consistent with the very low Ti content of late~stage magnetites (Table 7, analysis 8). The CaA12SiO6 content of groundmass clinopyroxene is lower than that

85.2 13.9 0.9

65.3 32.3 2.5

17.0 49,6 33.4

80.7 18.4 0.9

4.969

0.161 0.817 0.471

4.976

--

0.767 0.009 0.175

--

5.4 48.9 45,7

4.965

0.051 0.435 0.465

--

2.964 1.042 0,007

99.66

7,58 5.83

-1.06

65,86 19.65 0.19

T5 G

5

P = phenocrYst; C = core; G = gzoundmass; R = rim. A-B r e f e r t o c o e x i s t i n g alkali feldsIMur-plagiochtse pairs. a T h e s e p h a s e s c o , e x i s t w i t h s o d a l i t e in t h e g r o u n d m a s s .

An Ab Or

4.992

2.240 1.764 0.014

100.14

0.15 1.98

-15.70

49.13 32.82 0.37

T5 P,C

4

2,832 1;171 0.023

99.20

101.65

5.004

100.30

Total

5,49 5.37

-3.32

62.50 21.93 0.60

PS8 G

3

0.43 3.66

Total

0.15 1.50

K~O Na~O

0.19 13.41

2.360 1.614 0.029 0.013 0.644 0.024 0.318

0.28 16.63

MgO CaO

52.64 30.55 0.78

PS8 P,R

2

Atoms/8 oxygens: Si 2.179 AI 1.800 Fe 0,035 Mg 0.019 Ca 0.817 K O.009 Na 0.133

47.52 33.30 0.92

PS8 P,C

No.:

Wt.%: SiO~ AI20 ~ FeO

1

No.: Sample

Analysis

Representative feldspar analyses

TABLE 4

85.2 13.9 0.9

4.990

0.833 0.009 0.185

0.008

2.177 1.810 0.018

100.25

0.15 1.52

0.Ii 16.97

47.51 33.52 0.47

T27 P,C

6

64.1 33.0 2.9

4.959

0.595 0.027 0.306

0.007

2.407 1.598 0,013

99.91

0.47 3,49

0.II 12.28

53.22 29.98 0.35

T27 P

7

18.8 46.8 34.4

4.944

0.172 0.313 0.426

0.017

2.842 1,150 0,016

99.30

5,44 4,87

0.25 3.54

62.94 21.61 0.43

T27 P,R

8

84.7 15.3 0.0

4.977

2.196 1.801 0.014 . . 0.818 -0.148

100.88

-1.68

. 16.81

48.37 33.65 0.37

T6 P,C

9

.

44.4 53.2 2.4

4.975

2.572 1.439 0.006 . 0.426 0.028 0.510

99.54

0.40 5.85

. 8.84

57.16 27.13 0.16

T6 P,R

i0

.

6.7 81.4 12.0

4.974

0.064 0.115 0.985

2.942 1.068 --

100.41

2,06 9.23

. 1.37

67.08 20.67 .

T6 G,A

II

. .

--

0.0 30.4 69.6

4.966

-0.665 0.291

3.013 0.997

99.89

11.50 3.31

.

66.43 18.66 .

T6 G,B

12

6.92

4.977

0.331 0.044 0.588

2.664 1.350

99.96

0.77 6.80

34.4 61.1 4.5

.

59.76 25.70

T4 a G,A

13

1.9 29.1 69.0

4.966

0.018 0.657 0.277

2.987 1.028

99.58

11.31 3.14

0.37

65.62 19.15

T4 a G,B

14

cD

0.85

Mg/(Mg + Fe)

0.87

4.007

P = phenocryst; G = groundmass. aCore s a m p l e . b O u t e r zone s a m p l e .

4.003

Total

0.012 0.004

1.926 0.006 0.128 0.130 -0.912 0.890 .

.

99.91

100.18

0.76

4.016

-.

1.834 0.026 0.248 0,245 0,005 0,778 0.881

99.87

-.

-

0.16 0.15

-

0.16

0.14

--

14.01 22,06

49.23 0.91 5.66 7.85

3 PS8 pb

52.69 0.22 2.96 4.25

2 PS8 pa

16.74 22,74

-

15.89 23.14

-

52.06 0.38 3.67 4.89

1 PS8 pa

Atoms/6 oxygens: 1,905 Si 0.010 Ti 0.158 A1 0.150 Fe -Mn 0.867 Mg 0.908 Ca . . K -Na 0.005 Cr

Total

SiO~ TiO$ AI~O~ FeO MnO MEO CaO Na~O Cr203

Analysis No.: S a m p l e No.:

R e p r e s e n t a t i v e c l i n o p y r o x e n e analyses

TABLE 5

0.29

.

0.76

4.016

0.021

1.899 0.024 0.141 0.258 0,007 0.803 0.863

99.80

.

14.44 21.58

0.23

50.92 0.87 3.20 8.28

4 PS8 pb

.

.

0.73

4.014

0.041 .

1.943 0.017 0.097 0.286 0.013 0.770 0,845 0.004

98.85

0.56 .

13.68 20.89

0.39

51.46 0.58 2.18 9.04

5 PS8 G

.

0.73

4.011

0.012 . .

1.895 0.015 0.169 0,276 -0.750 0.896 --

99.44

0.16 .

13.41 22.27

--

50.48 0.53 3.81 8.78

6 T17 P

0.73

4.025

0.028 .

1,863 0.023 0.205 0.282 0.008 0.747 0.870 --

100.07

0.38 .

13.41 21.71

0.24

49.85 0.80 4.66 9.01

7 T5 P

0.66

4.024

0.057 .

1.919 0.028 0.124 0.352 0,011 0.666 0.859 0,008

100.0

0.78 .

11.83 21.21

0.36

50.79 0.97 2.78 11.12

8 T5 G

0.75

4.016

0.025

1,928 0.013 0.110 0.274 0.009 0.803 0.854 .

99.97

0.35

14.45 21.39

0.30

51.72 0.48 2.50 8.80

9 T13 P

.

0.76

4.035

0.033

1.858 0.020 0.206 0.238 0.011 0.766 0.904 .

99.03

0.45

13.65 22.39

0.34

49.32 0.70 4.64 7.54

10 T27 P

0.77

4.042

0.029 0.002

1,864 0.022 0.150 0.232 0.013 0.790 0.920 .

100.92

0.41 0.07

14.31 23.19

0.41

50.34 0.80 3.90 7.49

11 T32 P

0.73

4.016

0.016 --

1,914 0.015 0.148 0.268 0,008 0.727 0,898 .

99.89

0.53 --

13.05 22.41

0.26

51.20 0.54 3.35 8.56

12 T4 P

.

0.74

4.009

0.015 --

1.932 0.010 0.112 0.269 0.011 0.763 0.898

99.40

0.20 --

13.64 22.32

0.34

51.45 0.35 2.53 8.58

13 T6 P

0.73

4.024

0.028 --

1.879 0.020 0.181 0.270 -0.744 0.902

99.13

0.38 --

13.23 22.32

--

49.84 0.72 4.07 8.57

14 T28 P

12

Co

20 "*-- Mg

SO

I

I

40

50

X~" ~ c ~ c ~ r

9'o

~

~o

t 0 0 Mg/Mg + Fe

~o

rX

-

~o

ge do

20

l 30

t 90

810

I 50

40

~c~o-xo

i

70 60 OLIVINE

OLIVINE

50

×x-

r

50

~40

~x

40

30 ~( - °~0 80 OLIVINE

I

i

40

50

70

50

40 BIOTITE

6O

5O

CPX J

I

I

30

40

50

4tO

do

OLIVINE

;o

do

× 7b

~o

do

~o

CPX I 30 40

OLIVINE

do

;o

40

30

xQexr

70

/

CP/× I

do

~o BIOTITE 4,0

~o

I

41o

50

do

zb

~0

50

do

70

60

50

OLIVINE

£

BIOTITE .-1~4o

Fig. 3. A summary of the compositions of ferromagnesian minerals from Tambora lavas and intrusive rocks.

of the phenocrysts and decreases with decreasing Mg/(Mg + Fe) ratios. This is compatible with decreasing temperature and/or pressure (Fig. 4). One group o f phenocrysts from the most mafic trachybasalts (those from Satoncla island) have distinctly high Cr203 (0.15--1.0%) and high Mg/(Mg + Fe) ratios (0.85--0.9) (analyses 1 and 2, Table 5). These form clear, relatively unzoned cores to large phenocrysts with outer zones of more Fe-rich augite composition (Table 5, Figs. 3 and 4). Similar pyroxenes are described by Foden (1983) in ankaramites from Rindjani volcano and by Foden and Varne (1983) from other volcanoes in the east Sunda arc.

13

09 X

X

+ ~08

°°c

o o

[]

[]

o

on

ooo

[]

o (3

07 o

00t 0 02 003 Activity of CQAI2SiO 6 in CPX

004

Fig. 4. C o m p o s i t i o n s of c l i n o p y r o x e n e s f r o m T a m b o r a n e - t r a c h y b a s a l t s . Crosses are diopsidic cores f r o m large p h e n o c r y s t s , squares are p h e n o c r y s t s a n d circles are rims a n d g r o u n d m a s s . See T a b l e 13 for a c t i v i t y m o d e l used.

(3) Olivine The suite shows a silica undersaturated character in that olivine is present both as phenocrysts and in the groundmass. Representative analyses are given in Table 6 and are plotted in Fig. 3. Olivine phenocrysts typically have compositions in the range Fo~0--Fo65 and in general have slightly lower Mg/(Mg + Fe) values than coexisting clinopyroxene. This is particularly so in the coarse grained alkali gabbros where clinopyroxene with Mg/(Mg + Fe) values of a b o u t 0 . 7 - 0 . 7 5 coexist with olivine Foss. There is known to be a marked shift in the distribution of Fe2÷/Mg in favour of olivine coexisting with clinopyroxene at temperatures approaching the solidus in mafic intermediate systems (Powell and Powell, 1974; Parsons, 1981). This could support the contention that the coarse-grained rocks have cooled to lower temperatures than the lavas. Alternative explanations for the large difference in Mg/Fe 2÷ ratios of clinopyroxene and olivine in these rocks could include disequilibrium or magma mixing. In some of the more magnesian trachybasalts there are a range of phenocryst compositions (Fos0--FoT0) suggesting fractional crystallisation and re-mixing of segregated liquids and crystals. The trachybasalts from Satonda island have composite cognate xenoliths composed of olivine and clinopyroxene (+ magnetite) and these suggest that earlier fractionation is controlled by feldspar-free assemblages. Groundmass olivine is quite Fe-rich in the entire Tambora suite, ranging from Fo~0 to Fo40 with up to 2wt.% MnO.

(4) Magnetite The relatively high Fe2Oa content of the Tambora lavas (Table 2) has apparently resulted in a wide crystallisation interval of magnetite saturation. Titanmagnetite is a phenocryst phase in all lavas and a groundmass phase in

100.51

Total

3.006

0.64

Mn Mg Ca

Total

Mg/(Mg + F e )

0.72

3.009

0,989 0:004 0.571 0.014 1.429 0.004

100.33

37.39 0.11 25.81 0.60 36.28 0.13

2 PS2 P

0.56

2.997

1.001 0,005 0,861 0.030 1.088 0.012

100.26

35.70 0,15 36.72 1.27 26.03 0.39

3 PS2 G

0.81

3.009

0.989 0.004 0.389 0.008 1.611 0.007

100.76

39.03 0.14 18.33 0.38 42.62 0.27

4 T5 pa

P = phenocrysts, G = groundmass. aCon~ning melt inclusions with cpx, magnetite and glass. bContaining plagioclase--bearingmelt inclusions. cOlivine is mantled with biotite, table8, analysis 3.

Fe

0.994 -0.702 0.024 1.275 0,010

Si A1

Atorns/4 Oxygens

F e O

~0 MgO CaO

--

36.65

1 T13 P

30,93 1.06 31,52 0.36

SiO= A1203

Analysis No. : S a m p l e No.:

R e p r e s e n t a t i v e olivine analyses

TABLE 6

0.73

3.006

0.991 0,005 0.533 0.011 1.457 0.008

101.65

38.28 0.16 24.62 0.52 37.76 0.31

5 T5 pb

0.48

3.001

0.997 0.007 1.007 0.041 0.930 0.016

99.21

34.15 0.21 41.22 1.66 21.37 0.51

6 T5 G

0.74

3.002

0.998 -0.519 0.019 1.457 0.006

99.61

23.52 0.85 37.05 0.23

--

37.86

7 T32 P

0.42

2.993

1.005 0.004 1.113 0.065 0.806 --

100.93

45.44 2.61 18.46 --

0 . 1 2

34.30

8 T4 pc

0.50

3.002

0.998 -0.990 0.041 0.974 --

99.78

40.97 1.67 22.61 --

--

34.52

9 T4 p

0.54

3.010

0.989 0.905 0.029 1.078 0.010

100.84

38.44 1.20 25.70 0.33

--

35.16

10 T28 p

0.11

Mg/(Mg + F e 2÷)

0.08

0.44 0.05 0.42

24.000

3.384 0.847 -0.129 8.256 10.278 0.239 0.867

99.39

15.16 2.42 -0.54 36.96 41.40 0.95 1.96

2 PS2 G

0.07

0.62 0.09 0.26

24.000

2.069 1.854 0.042 0.048 9.916 9.244 0.131 0.694

98.13

9.22 5.27 0.18 0.2 44.15 37.03 0.52 1.56

3 T13 P

0.21

0.53 0.24 0.15

24.000

1.213 4.891 0.087 0.111 8.485 7.166 0.117 1.929

98.65

5.83 15.00 0.40 0.50 40.76 30.98 0.50 4.68

4 T5 P

0.10

0.56 0.11 0.29

24.000

2.288 1.810 0.094 0.170 9.350 9.060 0.172 1.056

100.18

10.48 5.29 0.41 0.73 42.81 37.32 0.7 2.44

5 T5 P

0.05

0.41 0.05 0.48

24.000

3.840 0.819 . 0.174 7.327 10.996 0.239 0.605

95.44

16.47 2.24 . 0.70 31.41 42.41 0.91 1.31

6 T5 G

.

.

0.19

0.54 0.13 0.21

24.000

1.679 2.050 . 0.113 10.480 7.702 0.156 1.822

99.86

7.78 6.06 . 0.49 48.54 32.10 0.64 4.26

7 T27 P

P = p h e n o c r y s t , G = g r o u n d m a s s , I = intrusive. F % O 3 calculated f r o m m i c r o p r o b e F e O values ( C a r m i c h a e l , 1967). 1. In m e l t inclusion in olivine, c o e x i s t i n g w i t h c p x .

0.48 0.08 0.35

24.000

XFe30~ XMgA1204 XFe2TiO *

Total

2.821 1.226 -0.189 8.942 9.499 0.186 1.136

100.13

Total

A toms~32 o x y g e n s : Ti A1 Cr V Fe 3+ Fe 2+ Mn Mg

12.84 3.56 -0.81 40.68 38.88 0.75 2.61

1 PS2 P

TiO 2 Al~O~ Cr203 V,O 3 Fe203 FeO MnO MgO

Analysis No. : Sample No.:

R e p r e s e n t a t i v e m a g n e t i t e analyses

TABLE 7

0.13

0.73 0.05 0.05

24.000

0.120 14.291 6.692 0.730 0.990

.

0.412 0.766

99.79

0.50 63.45 26.73 2.88 2.22

.

1.83 2.17

8 T6 I

0.03

0.58 0.04 0.30

24.000

-9.277 9.909 0.142 0.329

2.380 1.963

101.33

-42.49 40.84 0.58 0.76

10.91 5.74

9 T28 I

16

all holocrystalline lavas. The compositional trends of magnetite range from more Mg- and Al-rich, Ti-poor cores to Ti-rich rims with low A1 and Mg. In the intrusive rocks (Table 7) some late magnetite is almost pure Fe304 with very low Ti and this suggests that at the latest stages of evolution of the liquid, Ti may have been highly depleted as a result of magnetite and biotite precipitation. TABLE 8 R e p r e s e n t a t i v e b i o t i t e analyses Analysis No.: Sample No.:

1 T32 P

2 T32 P

3 T4 P

4 T4 P

5 T27 P

6 T6 I

7 T28 I

8 T28 I

SiO~ TiO~ Al203 FeO MnO MgO CaO Na20 K,O BaO C1

36.29 4.99 16.20 12.55 0.24 16.84 0.00 0.66 8.09 0.77 --

35.86 5.36 16.45 12.65 0.15 17.32 0.00 0.66 8.07 0.70 --

35.89 5.00 13.97 17.96 0.30 12,50 0.00 0.49 9.09 nd 0.06

36.74 3.86 14.90 15.43 -14.65 . 0.18 9.32 . --

35.52 5.50 15.31 11.07 -16.68

36.05 5.13 15.32 12.36 0.22 16.23 . 0.31 8.91 . --

35.57 0.27 14.44 20.95 0.22 13.00

36.06 0.50 14.45 17.50 -14.84

0.20 9.79

-9.30

0.11

--

Total

96.76

97.4

95.27

95.08

.

. 0.59 8.47

.

. -93.14

94.54

.

.

94.54

92.65

A toms~22 oxygens: Si Ti A1 Fe Mn Mg Ca Na K Ba CL Total Mg/(Mg + F e )

5.297 0.548 2.787 1.532 0.030 3.664 -0.187 1.506 0.044 --

5.202 0.585 2.813 1.535 0.018 3.746 -0.186 1.494 0.040 --

5.476 5.518 5.335 0.574 0.486 0.621 2.512 2.637 2.710 2.292 1.939 1.390 0.039 --2.843 3.278 3.734 0.000 . . . 0.144 0.053 0.171 1.770 1.785 1.623 . . . . 0.015 ---

5.374 5.566 0.575 0.031 2.691 2.663 1.541 2.741 0.028 0.029 3.605 3.033 . . 0.090 0.062 1.694 1.955 . . -0.029

5.622 0.059 2.655 2.282 -3.449

15.609

15.639

15.665

15.646

15.585

15.597

16.108

15.916

0.71

0.71

0.55

0.63

0.73

0.70

0.53

0.60

-1.849 --

P = p h e n o c r y s t ; I = intrusive. 1. Ba is o n l y d e t e r m i n e d o n analyses 1 and 2 a n d m a y be p r e s e n t in t h e o t h e r a n a l y s e s as well. 2. Biotite r i m m i n g olivine, analysis 8, Table 5.

17

(5) Biotite Biotite occurs as phenocrysts in the glassy trachyandesites and both as an early phenocryst phase and later as a reaction product in the intrusive rocks (Fig. 3). In the latter role it occurs as rims around magnetite and olivine. In the trachyandesites, biotite is moderately magnesian with Mg/(Mg + Fe) ratios in the range 0 . 7 - 0 . 7 5 . These have high A1 content, high Ti (5.0--5.5 wt.% TiO2), high Ba and are chlorine-bearing (Table 8). The phenocrysts of biotite in the intrusive rocks are of the same composition as those from the trachyandesites b u t those forming reaction rims are much more Fe-rich with Mg/(Mg + Fe) values of 0.5--0.65. The most Fe-rich of these are quite Tipoor.

(6) Other phases The remaining important phases are leucite, apatite and sodalite. Apatite is c h l o r o h y d r o x y apatite and sodalite appears to be a pure chlorine end member (Table 10) and these together with the presence of C1 in the biotite provide evidence of significant activities of chlorine in the melt. Apatite is present as a groundmass phase in the more magnesian trachybasalts and as a microphenocryst phase in the low Mg trachybasalts, trachyandesites and inTABLE 9 Representative leucite analyses Analysis No. : Sample No. :

SiO~ Al~O 3 FeO CaO K20 Na~O Total

1 T30 MP 55.47 21.84 -1.2 21.49 -100.50

2 T30 MP

3 T30 G

54.50 21.84 0.15 1.11 22.30 --

54.7 21.67 -1.18 22.4 --

99.9

99.95

Atoms~6 oxygens: Si A1 Fe Ca K Na

2.022 0.939 -0.047 0.999 --

2.005 0.946 0.005 0.043 1.047 --

2.010 0.938 -0.046 1.051 --

Total

4.007

4.046

4.045

MP = m i c r o p h e n o c r y s t ; G = g r o u n d m a s s .

18 T A B L E 10 R e p r e s e n t a t i v e s o d a l i t e analyses A n a l y s i s No.: S a m p l e No. :

1 T4

SiO 2 A120 ~ K20 Na20 C1

37.18 31.54 0.72 23.27 6.83

38.02 32.02 0.71 22.33 7.26

37.66 31.65 0.67 24.04 7.15

Total

99.54

100.34

101.17

1.54

1.64

1.61

O --- C L Total

98.00

2 T4

3 T4

98.7

Sodalite intergrown with K-feldspar, analysis 14, Table 4. T A B L E 11 Microprobe analyses of glass from ne-trachyandesites T32

T27

SiO 2 TiO 2 Al20 ~ FeO MnO MgO CaO Na20 K20 P2Os

58.35 0.52 19.34 4.5 0.13 1.21 3.39 5.02 6.75 0.28

58.5 0.58 19.70 4.20 -1.4 3.98 5.4 5.9 0.24

Total

99.49

99.9

CIPW n o r m : or ab an ne di ol mt il ap

39.89 34.02 10.30 4.58 3.79 1.88 3.55 0.99 0.65

34.87 36.99 12.09 4.71 4.87 1.64 3.33 1.10 0.56

Q Ks Ne

41.8 28.9 29.3

41.8 25.9 32.3

99.56

19 trusive rocks. Sodalite was only recognised in some of the intrusive rocks (T6) and is a late-stage crystallisation product intergrown in a eutectic fashion with K-feldspar. Representative analyses are given in Table 11. It should be noted that most mineral analyses were carried o u t using the KEVEX energy
20

The compatible elements (Ni, Sc, Cr, V, MgO) show continuous depletion trends from MgO-rich trachybasalts to MgO-poor trachyandesites (Fig. 6). Conversely incompatible elements of the group K~O, Ba, Zr, Nb, and Rb show fairly linear trends of continuous enrichment (Fig. 5). This implies that the entire group of Tambora rocks analysed may belong to a common suite and that the divergent trends involving CaO, Sr and A1203 do not indicate separate unrelated magma types. This is also supported by the identical STSr/86Sr ratios of trachybasalts and trachyandesites (0.70385--0.70399) (Whitford et al., 1978). The difference between the trachybasalt and trachyandesite trends can be explained as the result of differences in the proportions of plagioclase in the crystallising assemblage. In the trachybasalt group the ratio of clinopyroxene and olivine to plagioclase could be high resulting in a bulk distribution coefficient for Sr considerably less than one. CaO depletion would result from clinopyroxene removal. In the trachyandesite series the proportion of plagioclase is much greater and both Sr and A1203 are depleted as a result. As a test for the above model a least-squares mixing approach was adopted (e.g. Bryan et al., 1969). The results of these calculations are presented in Table 12. Those results given in Table 12a model the derivation of low MgO trachybasalt T5 from high MgO trachybasalt T12. This is a model for differentiation in the Tambora magmatic system prior to 1815. The predicted cumulate is a clinopyroxene rich gabbro with a high clinopyroxene and olivine: plagioclase ratio (2:1). 8

4o

6

30

~4

E 20

2

IO

o

0

÷

X

,,'o

8'0

,2o'

I

~o

x

x 0

I

I

2

2OO

ppm Rb

o

I

4

Wl%

6

MgO

50

BL

40

~

30

iI

~ 20

x

×x

a

I

i

8O0

ppm 80

x

[3 XQ~]+ ~0

I0 I

,too

0

~4oo

Wt°lo MgO

Fig. 5. Whole rock analyses of Tambora lavas and intrusive rocks, S y m b o l s are the same in Figs. 5, 6 and 7. Squares = high MgO ne-trachybasaits; c~osses ffi low MgO ne-trachybasaltsi circles = ne-trachyandesites ( 1 8 1 5 eruption); plus signs = intrusive rocks (syenites). Fig. 6. C o m p a t i b l e trace e l e m e n t versus MgO variation for Tambora lavas. S y m b o l s as in Fig. 5.

21 TABLE 12 Least-squares model testing of possible crystal--liquid relations in the evolution of Tamborn lavas (a) Least-squares approximation to trachybasalt T12 in terms of low MgO trachybasalt T5 and phenocrysts

SiO~ A1203 FeO MnO MgO CaO Na20 K20 TiO~ P205

T12 Est.

T12 Obs.

Cumulate composition Component

48.29 17.93 10.13 0.18 5.24 10.42 3.45 2.58

48.26 17.95 10.12 0.21 5.27 10.46 3.50 2.67

40.69 11.19 19.76 0.29 10.54 14.58 0.95 0.04

0.91 0.47

0.97 0.42

1.53 0.44

Wt. fraction

T5 (liquid) 0.7224 cpx. 0.1255 plag. 1 0.0754 mag. 0.0482 olivine (Fo 73) 0.0269 apatite 0.0028

Cumulate wt.%

45.0 27.04 17.29 9.65 1.00

% crystallisation = 27.8 (wt.) ~ residuals2 = 0.0211

(b) Least-squares approximation to low MgO trachybasalt T30 in terms of trachyandesite T23 and phenocrysts T30 EST

T30 OBS

Cumulate composition Component

Wt. fraction

Cumulate wt.%

SiO 2 A120 ~ FeO MnO MgO CaO Na~O K~O

51.49 20.54 6.52 0.14 3.18 8.71 3.83 3.29

51.39 20.56 6.52 0.15 3.25 8.89 4.42 3.56

47.29 21.96 7.51 0.09 4.18 13.46 2.73 0.85

0.5785 0.0712 0.2712 0.0401 0.0231 0.0095

17.15 65.33 9.66 5.56 2.29

TiO 2 P2Os

0.79 0.71

0.68 0.49

0.99 0.94

T23 (liquid) cpx plag 2 biotite mag. apatite

% crystallisation = 42 2: residuals 2 = 0.5320

(c) Least-squares approximation to trachybasalt T12 in terms of trachyandesite T23 and phenocrysts

SiO2 A1203 FeO MnO MgO CaO Na20 K~O TiO 2 P205

T12 EST

T12 OBS

cumulate composition Component

Wt. fraction

Cumulate wt.%

48.33 17.93 10.11 0.16 5.20 10.34 3.11 2.46

48.26 17.95 10.12 0.21 5.27 10.46 3.50 2.67

45.28 17.14 12.34 0.15 6.62 12.93 2.35 1.15

0.3347 0.1989 0.3171 0.0697 0.0675 0.0071 0.0093

29.7 47.4 10.4 10.1 1.06 1.39

1.17 0.57

0.97 0.42

1.44 0.59

T23 cpx plag 3 biotite mag. Olivine apatite

% crystallisation = 66.5 (wt.) residuals 2 --- 0.288 (continued)

22

(d)

Variable compositions cpx

plag 1

plag

CaO Na20

50.27 3.71 8.64 0.17 13.42 22.04 0.20

50.83 30.65 0.83 . . . -14.32 3.13

53.60 29.63 0.90 . . . -12.93 4.03

K~O

--

0.I0

0.I0

SiO~ Al~O~ FeO

MnO MgO

TiO~ P2Os

0.57

.

.

.

.

2

. .

.

plag 3

olivine

mag.

biotite

apatite

T23

53,87 28.46 0.76

37.90 -23.50 0.85 37.05 0.23 . . . .

--

36.29 16.20 12.55 0.24 18.84 -0.66

-----52.40 --

55.17 19.78 5.9 0.18 2.50 5.42 4.66

--

--

-11.53 4.61 0.61

6.19 75.50 0.67 4.33 --

.

7.17 .

.

8.09

--

5.08

4.99

-41.50

0.67 0.55

.

Results in Table 12b illustrate a test of the possibility that liquids differentiate from parent high MgO trachybasalt to trachyandesite via low MgO trachybasalt. The residuals produced by these calculations are quite high. A second calculation (Table 12c) only considers the relationship between parental high MgO trachybasalt and trachyandesite and is therefore independent of the path taken by the liquid. The residuals are lower in this calculation than in Table 12b. The results of calculations, presented in Tables 12b and c both indicate that the ratio of clinopyroxene to plagioclase is less than one, which accords with the inference drawn earlier. In Fig. 7 the 1815 trachyandesite group is separated from the trachybasalt £4

,2 I

~2c :4

c~

J6 +

5oo

~o

~o

' HOG

,~o

J2

i~oo

500

~oo

ppm Sr

,l~O

1300 '

If~O

ppm Sr

~;

E 8OO F

~o

.~ o~

~

& I

~oo

~oo

~

,,oo-,~oo p p m Sr

F

,5°°

l

. . . .

;--; WI%

~; AI203

22 2

F i g . 7. C a O , ,4_1203, Sr a n d Ba interrelationships amongst Tambora lavas and intrusive rocks. Symbols as in Fig. 5. The arrows illustrate the different liquid evolutionary trends for: (1) the generation of low MgO trachyba~alts from high MgO trachybasalts; and (2) the generation of trachyandesites (1815 products) from trachybasalts.

23 trend which also extends to considerably higher Sr values (see CaO v Sr diagram, Fig. 7). This suggests t h a t these most St-rich, fractionated trachybasalt liquids do not further fractionate to yield the trachyandesites. This separation of the trachyandesite group suggests t h a t the 1815 eruption disgorged a magma chamber whose contents had evolved as a closed system. In this case the liquid evolutionary path m a y have been akin to equilibrium crystaUisation. It is also possible that the coarse-grained rocks (e.g. T6, Tables 2 and 3) represent side-wall or r o o f crystallisation in this magma chamber. This contrasts with the trachybasalt trend which is continuous and indicates open system fractionation where batches of melt are regularly extracted and erupted t h r o u g h o u t the course of crystallisation. Figure 8 illustrates the main features of the evolution of the Tambora lavas as discussed in this section. ~'C PX

CnO

TRACHYBASALTS Open system O ~ o n o I crystoIlizotion

.......

l

A TRACHYA~

PLAG

/

creos, fir- PLAG/CPX

rOtiO

Closed system Reslduolliquid after 50-70% Equilibrium crystoPhzotion CPX / PLAG ~-- 05

Sr or AI203

'p

Fig. 8. Schematic diagram illustrating the separate evolutionary paths taken by Tambora liquids with the relatively clinopyroxene-dominated trachybasalt path and the separate, discontinuous trachyandesite (1815) liquids generated in a closed system by increasingly plagioclase-dominated precipitation. CONDITIONS OF ERUPTION Final eruption of the system was described as having begun with the emission of some dark clouds and with explosions on 5 April 1815 (Stothers, 1984). Stothers (1984) and others (Self et al., 1984) suggest that this initial activity followed a few days of rumbling as the only warning of any impending eruption. The eruption reached its paroxysm on 10 April and the explosion is reported to have been heard as far as 2600 km away (Stothers, 1984). Many areas within 600 km of the volcano are reported to have remained in total darkness for 1--2 days after the eruption and the seas to the north of Tambora were the site of large pumice rafts, some several kilometers long. The eruption column probably penetrated the stratosphere to an altitude of more than 17 km (Stothers, 1984) and the dust, aerosols and water vapour (ice) were then globally distributed producing significant climatic effects, modified sunsets and unusual twilights of between one and four years duration (Self et al., 1984).

24

Volume of ejecta Hammer et al. (1980) have correlated acid impurities in Greenland ice cores in the years 1815 to 1818 with fallout from the T a m b o r a eruption and have calculated that global fallout of acid aerosols attributable to this event is of the order of 1.5 × 1014 g. If the relationship between the concentration of these volatile c o m p o n e n t s in the magma and the extent they are exsolved to the atmosphere is known, then Hammer et al.'s (1980) estimate of the total atmospheric load o f acid aerosol could be used to estimate the weight of magma discharged in the 1815 eruption. Some C1 and S analyses are reported in Table 3. The 1815 lavas have 3 to 4 times the concentrations of these elements in the trachybasalts. Devine et al. {1984) also report the F content of the 1815 trachyandesite as 1185 ppm. The solubility of sulphur in silicate melts is reported by Haughton et al. (1974) and is a complex function of T, fs2, fo~ and melt composition. The solubility of chlorine in magmas has also been investigated b y Kilinc and Burnham (1972). Devine et al. (1984) have studied glassy eruptive products from a large n u m b e r of historically active volcanoes. They have determined S, C1 and F contents of coexisting groundmass glass and glassy inclusions in phenocrysts to estimate the loss of these volatiles on eruption. Tambora was included in this study. Devine et al. {1984) calculated the following yields during the 1815 eruption of Tambora; F = 506 ppm, C1 = 895 p p m and S -- 72 ppm. These convert to 533 ppm HF, 920 p p m HC1 and 220 ppm H2SO4 giving a total of 1673 ppm acid aerosol discharged to the atmosphere. Using Hammer et al.'s (1980) estimate of fallout (1.5 × 1014 g) this yields a total weight of magma of 8.97 × 101~ g or 33.2 krn 3 at a density of 2.7 g cm -3. Published estimates of the volume of material erupted by the 1815 blast vary considerably. The most recent and best is that of Self et al. (1984). They estimate that ash fall and ignimbrite deposits total a b o u t 200 km 3 distributed over a density range from 600 kg m -3 to 2700 kg m -3. This reduces to close to 50 km 3 of dense rock and magma. The caldera has a volume of a b o u t 36 km 3 (6.5 km in diameter and between 1000 and 1200 m in depth). This together with the destroyed 15 km 3 of the upper cone (Self et al., 1984) yields a b o u t 50 km 3, the same as the estimate based on the pyroclastic deposits. The volume of magma estimated using the aerosol fallout (33 km 3) is very close to the caldera volume which may be expected to be mainly created by the disgorgement of the magma chamber. Again the addition of the 15-km 3 host portion of the pre-1815 volcano yields a total volume of displaced rock and magma of a b o u t 50 km 3. The energy rleased by the eruption has been estimated by a number of workers as close to 102v ergs ( Y o k o y a m a , 1957; Allard et al., 1983). Such estimates are prone to large errors as they are based on the estimation of

25 ejecta volume and of the proportion of that ejecta that started the eruption at magmatic temperatures. The estimate presented here suggests that a b o u t 70% of the mass of ejecta was at magmatic temperature. The principal energetic source of an eruption is the heat energy (temperature, specific heat and heat of crystallisation) of the magma. This is estimated as a b o u t 1.4 -+ 0.3 × 10 l° ergs g-1 of magma {Williams and McBirney, 1979). On this basis the estimated 8.9 × 1016 g of magma could supply about 1.2 × 1027 ergs equivalent to a b o u t 30,000 megatons of TNT. Discharge velocity

Several accounts (e.g. Petroeschevsky, 1949) refer to the ejection of bombs and rock fragments ("the size of t w o fists") a distance of as far as 40 km from the summit of the volcano. The range of such fragments if these are assumed to have travelled a ballistic trajectory is as follows: R = V 2 sin 20/g Where V is the initial velocity, g the acceleration due to gravity, R the range and 0 the ejection angle, neglecting the effects of wind resistance. If the debris observed at 40 km from the blast has travelled a m a x i m u m distance then the above equation becomes (0 = 45°): R = V2/g

For R = 40 km this gives a discharge velocity of 630 m s- 1. This is very similar to estimates of discharge velocities from other recent very large volcanic blasts, for example the 1968 eruption of Arenal volcano in Central America (Fudali and Melson, 1971). An eruptive pressure can be calculated from this projectile velocity based on the relationship: w pA =g

dv -dt

where P is pressure, w the weight of the projectile and d v / d t its acceleration. This has been reduced to P = 0 . 0 1 2 5 V 2 (Minakami, 1950; Decker and Hadikusumo, 1961) when appropriate constants are substituted for volcanic conditions, yielding an estimate of the discharge pressure (P) of 4898 bar. Again this is very similar to values calculated by other workers for very large blasts. Such a discharge pressure (nearly 5kbar) is obviously very high and it is interesting to try to understand how such a vast overpressure of vapour was generated. The following section approaches this problem from the petrological standpoint, both to independently check the pressures calculated above and to generate a model for the magmatic history to the blast.

26

Temperature, pressure and oxygen fugacity Trachybasalts and low-MgO trachybasalts The compositional variation of the trachybasalts has been attributed to fractional crystallisation. In the Fo-Q-Ks system (Fig. 9) the relative proportion of olivine with respect to clinopyroxene decreases through the series in accordance with the expected down-temperature curvature of this cotectic (e.g. Irvine, 1970). The compositions of magnetite phenocrysts also provide evidence of cooling during crystallisation. Cores have significantly higher Mg and A1 contents indicating higher MgA1204 (spinel) contents. Rims have higher titan -- magnetite contents (Table 7). This is indicative of the general trend where pyroxene and olivine are formed at the expense of aluminous spinel components as temperature and/or pressure decrease. Such a reaction for example could be represented; MgAI204 + CaMgSi206 ~ Mg:SiO4 + CRAI2SiO6 spinel

diopside

olivine

(1)

CATS--cpx

(Herzberg and Chapman, 1976). Ferromagnesian phenocrysts in general show a range of Mg/(Mg + Fe) values suggestive of changing liquid compositions and individual samples often contain more than one population of the same phenocryst type. For instance trachybasalt PS 2 has two sets of olivine phenocrysts: one with compositions of FoB0 and the other about FoT0. The (a)

Fo

Lc

(b)

Q

Sa

Fig. 9a. Tambora rock compositions plotted in the system Fo-Ks-Q with the fields of leucite (Lc) and phtogopite (Phi) from Luth (1967)superimposed. Fields with hatchuring and dashed margins are at PH~O ffi 5 0 0 0 barl O p e n fields defined by solid lines are at PH~O = 1000 bar. The position of a typiea! biotite from the 1815 trachyandesites is shown. Curvature of the t r a c h y ~ l t - d e f i n e d trend suggests the increasing rote Of pyroxene crystallisation relative to olivine. Fig. 9b. Composition of Tambora rocks plotted in the system Ne-Q-Ks. Symbols as in Fig. 5.

27

former contain melt inclusions containing clinopyroxene, glass and magnetite, the latter have inclusions containing plagioclase, clinopyroxene, magnetite and glass. The phase relationships implied by both the change in composition of these phenocrysts and by the appearance of plagioclase is clearly indicative of decreasing temperature. Two reactions which are functions of temperature and silica activity are: CaA12SiO6 + SiO2 ~ CaA12Si208 CATS-cpx

-ln(asio2 ) =

melt

(2)

plagioclase

5 9 , 6 4 0 - 4.1 T - 1.002 P 8.3144 T

-

ln(aAn) + ln(aCATS)

(Arculus and Wills, 1981) and KA1Si206 + SiO2 ~ KA1Si~O8 leucite

log(asio~) =

melt

(3)

sanidine

(-972 T

0.039 -

0.0345 (P T

17,)/ + log(asan) -log(aleucite) J

(Nicholls et al., 1971). Leucite however is generally only a groundmass phase or occasionally a microphenocryst phase in the trachybasalts. Empirically, this absence of leucite phenocrysts given the position of these rocks in the Q-Ne-Ks system suggests that H~O has contracted the leucite field relative to its 1 atmosphere position (Scarfe et al., 1966). This is particularly so for the trachyandesites. Clinopyroxene is present continuously through the crystallisation interval and reaction (2) predicts that clinopyroxene in equilibrium with liquids of a given asio2 will have higher activities of CaA12SiO6 at higher T and P. Thus groundmass clinopyroxene should perhaps show trends of decreasing activity of calcium tschermak molecule (CATS) compared with values from phenocrysts. This tendency is indeed illustrated by the analyses presented in Fig. 4. This tendency is also likely to be enhanced by the large effects of increased activity of H20 in displacing the composition of plagioclase to higher An contents (Yoder, 1969; Johannes, 1978). This would effectively displace reaction (2) to the left. In Fig. 4 a group of highly magnesian, chrome diopside compositions from the cores of some large, zoned clinopyroxenes (Table 5, analyses 1 and 2) have very low CATS contents. Such pyroxenes have been noted from a number of other island arc suites (e.g. Rindjani volcano, Lombok; Foden, 1983) and Foden and Varne (1983) suggested that they are xenocrystal, perhaps of lower crustal origin, later overgrown by new clinopyroxene from the liquids in which they were entrained and carried to the surface.

28

The solutions to eqs. (2) and (3) in terms of temperature and asio2 (melt) are plotted in Fig. 10 for the trachybasalts and low Mg trachybasalts. The activity data used in these and future calculations are given in Table 14. These were calculated from electron microprobe analytical data using the activity models given in Table 13. The results plotted in Fig. 10 suggest groundmass quench temperatures of the order of 850--900°C and asio: s5

/',:>0 ) b0r,,

:J~,:J~

*

:~r :3o, r~

I/

C'a

GSI0203 reel! 02

I

melt ~ !

"I--:'2;o

01' o

/

I

Jl

i

_-::,rz~'

//(/

....

.... : t....

T °C

T°C

Fig. i0. Curves marked 2 are solutions to eqn. (2) and curves marked 1 are solutions to eqn. (3). Solutions are in terms of T and asio~ (melt) at specified pressures. The lefthand figure shows results from high MgO-trachybasalts while the right-hand shows results from low MgO-trachybasalt and trachyandesite. Solid lines denote results on groundmass mineral assemblages, dashed lines on phenocryst assemblages (which do not include leucitebearing assemblages). Component activities used are given in Table 14 calculated according to models in Table 13.

TABLE 13 Components and activity expressions used in calculations Component

Activity

Reference

CaAl,Si20,

(X~a)

Kerrick and Darken (1975)

NaAISi~Os

(X~Na), (X~Na)

Kerrick and Darken (1975)

KAISi30 s

(X~K ), (XK ~)

Kerrick and Darken (1975)

F%SiO4

(~Fle)

Kerrick and Darken (1975)

Mg2SiO4

(X°kg)

Kerrick and Darken (1975)

CaAl2SiO,

( X~a ), ( XM~ )

Herzberg (1978)

KA1Si206

(XI~)

Nicholls et al. (1971)

F%O~

X Ft%0"

Carmichael et aL (1974)

K~F%(Si6AI2020)(OH)4

(xb~'~t )

Wones and Eupter (1966)

T4 core rim interstices T6 core rim interstices

5. intrusives

0.80 0.59 0.34 0.85 0.45 0.7

0.84 0.74

0.85 0.65 0.12

0.85 0.80 0.60 0.53

aPl an

0.20 0.39 0.60 0.15 0.55 0.81

0.16 0.25

0.15 0.33 0.69

0.15 0.2 0.38 0.44

a~ab

0.12

0.24

0.30

0.01

0.02 0.19

0.02 0.03

aPl or

0.02 0.06

0.50

0.49

0.50

aaf ab

0.70

0.74

0.34

0.46

0.38

aaf or

0.292

0.073

0.281

0.076

0.122

0.081

afa ol

0.007

0.01

0.015

0.006

0.016

0.006

0.014

aCpX cats

0.98

0.98

aK lc

0.45 0.30

0.37

0.265

abiot ann

0.73

0.5

0.41

0.56

0.435

0.535

a Ti'mt mt

264

258

249

285

1.

274

281

286

292

2.

1. Me(X m < 0.5) where Me = albite-equivalent mass of melt of the given c o m p o s i t i o n w h e r e the m o l e fraction of H 2 0 is a) ~< 0.5 (Burnham, 1978). 2. Me(Xw m > 0.5) where the mole fraction of H20 is > 0.5.

T32, analysis 1 table 11

(b) groundmass

(a) phenocrysts

(b) groundmass

(a) phenocrysts

(b) groundmass

(a) phenocrysts

4. glass

3. ne-trachyandesite (T27, T32)

2. ne-trachy basalts (T30, T5)

low Mg

1. ne-trachy basalts (T12, T17, T8)

C o m p o n e n t activities used in calculations

T A B L E 14

t~

30 (melt) values of close to 0.3. Using these values it is possible to evaluate fo: by using the reaction: 3Fe~SiO4 + O2 ~ 2Fe304 + 3SiO2 olivine

gas

(4)

magnetite melt

~ log f% = -- 2-5 5T8 2 + 8.904 + 0.13 T( P - 1) _ log ~"Fa' t,,OL~3 + log ~,Fe304 /,,Mt ~2 -

log (aSiO2) 3

(Nicholls, 1977; Arculus and Wills, 1981). The results of these calculations are plotted in Fig. 11. They suggest log fO, values at the temperatures calculated above o f - 9 . 5 to - 1 0 . 5 , which are quite high ( a b o u t two to three log units below the magnetite-haematite buffer). These values were also checked using the model of Sack et al. (1980) based on the FeO/Fe203 ratio. The log fo~ values from this m e t h o d range from - 1 1 . 4 7 at 800°C to - 7 . 6 3 at 1000°C (Fig. 11) and as such are very close to those arrived at before. The estimation of phenocryst equilibrium temperatures is difficult as the assemblages do n o t provide reliable geothermometers. Plagioclase phenocrysts are highly zoned and their cores frequently are strongly corroded. Under these circumstances co-existing plagioctase-melt pairs cannot be reliably identified. Leucite is n o t present as a phenocryst phase and therefore the approach adopted above cannot be ufilised. The asio, - T curve for reaction (2) utilizing clinopyroxene and plagioclase core compositions at 2000 bar .... j

-8

j -4

f

-6

~

Log -8

~

~

~

~

~

Le., -14

j l

J-~J

~

. .J

@

/1/

~enocrysTol

j~'~

J

./

7(~0°

~oo°

S ~'" 1

,~°

Temperature

,~°

,,~o.

°C

Fig. ll. Variation of logfo2 with temperature. Curve I (trachyandesite)was calculated from (4) for a pressure of 5000 bar and with aslo-T pairs from Fig. I0. Curve 2 shows the range of fo2 values calculated using the metho~ of Sack et al. (1980) for the trachybasalt whole-ro~k compositions and analysed FeO/Fe20 s ratios at 1 atmosphere, Individual points are calculated from eqn. (4) and utilize asio, and T pairs determined in Fig. ~0. The T17 phenocryst result assumed a temperature of ~050°C, P ffi2kbar and asio~ -- 0.29. Activities of magnetite and fayalite used are given in Table 14.

31 is shown on Fig. 10. If asio2 at this stage is close to that on eruption then this suggests phenocryst equilibration temperatures !50--200°C greater than those of the quench temperatures. Herzberg and Chapman (1976) have calibrated reaction (1) in spinel lherzolite assemblages at 12 and 16 kbar. They determined at 16 kbar that for reaction (1) in K = - 9 2 3 1 / T + 4.43 where: K = acpx CaA12SiO6 a°l Mg'SiO4 acpx CaMgSi20~ asp MgAl~O4 The effects of pressure on this system m a y be large and at the much lower pressures in question here and with very low aMgA1204 in magnetite solid solution, the specific temperature values derived m a y n o t be very meaningful. However the temperature difference indicated' by the phenocryst and groundmass assemblages is more likely to be meaningful. The equilibrium constant K for reaction (1) was determined for the phenocryst assemblages as 0.1979 (aMg2SiO' = 0.64, acaAhSiO6 = 0.016, aMgAl~O4 = 0.09, acaMgSi~O~ = 0.575) and at 0.0874 for the groundmass (aMg, SiO. = 0.42, aCaAl~SiO~ = 0.006, acaMgSi O = 0.58 and aMgA10 = 0.05). At 1~ kbar these would indmate temperatures of 1253 C and 1071 C or a temperature difference of 182°C. If, as calculated earlier quench temperatures are of the order of 900°C then this suggests phenocryst temperatures of a b o u t 1080°C. The co-existing groundmass plagioclase-alkali feldspar compositions also provide an estimate of quench temperature using the m e t h o d of Stormer and Whitney (1975). Such calculations for the low Mg trachybasalts (e.g. T30) yield temperatures very close to those already estimated (900°C, Table 15). The groundmass feldspars of the high Mg trachybasalts all contain substantial anorthite components (Fig. 2) and derived temperatures are geologically unreasonable. •

2

6

0

2

4

0

•

•

Trachyandesite These products of the 1815 eruption are glass-rich and show evidence of elevated activity of H20 in a number of their characteristics. As already discussed they contain no leucite yet fall well within the 1 bar leucite field in the system Q-Ne-Ks (Fig. 9). Similarly on the Fo-Ks-Q ternary system (Luth, 1967) liquids of these compositions would fall well into the 1 bar leucite field (Fig. 9). However, as determined by Luth (1967) the field of phlogopite (biotite) extends between the fields of olivine and leucite as the PH20 increases, which is very well illustrated by the mineralogy of the trachyandesites (olivine + An-rich plagioclase + clinopyroxene + magnetite + biotite + glass). Sanidine is mainly formed during quenching and is present as thin rims on plagioclase and as small microlites. In the absence of leucite, temperatures could only be defined as a function of asio~ over a range of values of temperature. These T - asio~ pairs (Fig. 10) were then used in reaction (4) to define a range of fo~ values as a function of T. These are plotted in Fig. 11 and values are one to two log

32

units higher than those of the trachybasalts. The calculations presented previously suggest that these rocks represent liquids formed by some 70% crystallisation of the trachybasalt parent liquids. Evidence already presented suggests the phenocrysts of these trachybasalts equilibrated at about 1080°C and empirically 70% crystallisation must require at least 100°C of cooling, suggesting that the trachyandesites were erupted at 900°C or less. Intrusive rocks

The intrusive rocks have a mineralogy suggesting crystallisation in a closed system. They are not cumulates and have whole-rock compositions which do not distinguish them very markedly from the lavas (Table 2). Most have clinopyroxene, magnetite and plagioclase compositions like those of the phenocrysts of the lavas. These phases then show evidence of reaction at lower temperatures with liquids of high alkali content and perhaps higher fo2. Thus magnetite and olivine are partially reacted to biotite, new low Ti magnetite is formed and alkali feldspar, sodic plagioclase and sodalite fill the interstices. These features suggest this group of rocks represent mixtures of accumulated phenocrysts with large amounts of evolved intercumulate melt. Coexisting feldspar geothermometry (Stormer and Whitney, 1975} suggests temperatures of the order of 700°C at PH20 = 4--6 kbar (Table 15). This temperature must then represent the lowest range of temperatures in this magmatic system and on textural evidence is that of the solidus. T A B L E 15 Alkali feldspar--plagioclase t h e r m o m e t r y ~ Sample no

Pressure ( b a r )

T e m p e r a t u r e °C

T30

1 1000

905 922

T6

1 1000 4000 6000

597 610 648 674

T4

1 1000 4000 6000

668 681 723 750

i Using m e t h o d o f S t o r m e r a n d W h i t n e y ( 1 9 7 5 ) .

Water vapour pressures

The existence of a hydrous phase, biotite in the trachyandeBitem, allows the estimation of PH~O ~ u e s at the time of eruption of Tambora: Composi~ tions of biotite are given in Table 8, Microprobe analyses show that there

33

is a significant C1 c o m p o n e n t in these biotites and this must reduce the activity of h y d r o x y biotite end members calculated. Wones and Eugster (1965) and Wones (1972) have calibrated the free-energy change o f the reaction: KFe3A1Si3010(OH)2 + 02 X KA1Si3Os + Fe304 + H20 aanite

gas

sanidine

mag.

(5)

steam

They determined that: log f H 2 0 =

3 4 2 8 - 4212 (1 - X) 2 T + l o g X + l o g f o : + 8.23 - l o g a s a n - log aMt

Where X is the mole fraction of annite in h y d r o x y biotite. This value will be less than the mole fraction Fe/(Fe + Mg) due to the presence of oxybiotite and chlorobiotite. Table 16 shows the range of fH20 values obtained for the trachyandesites using fo~ and T pairs on the buffer curve in Fig. 11. Two sets of values are presented, the first calculated by taking activity of annite to equal total Fe/(total Fe + Mg) (a maximum) and the second by taking the activity of annite arbitrarily less than this at 0.1, to take account of factors discussed above. Water fugacities thus obtained were then converted to PI-I:O values using the fugacity tables of Burnham et al. (1969). It is clear from the results of these calculations, that the fugacities of H20 indicated increase very rapidly as higher temperatures are chosen. Without a precise estimate of the temperatures of the trachyandesites an e x a c t fH~O value cannot be arrived at. However, it is very likely that these magmas were between 700°C and 900°C at the time of eruption and given the uncertainties in the activity of annite this allows theoretical fH~O values to reach astronomical levels. It is clear however that the data can easily support the

T A B L E 16 Water f u g a c i t y / p r e s s u r e of T a m b o r a t r a c h y a n d e s i t e m a g m a I A. aannite = 0 . 2 6 5 = m o l e f r a c t i o n of a n n i t e logfo. - 12.95 T (°(~ 700 fH, o ( b a r ) 1356 p~20 2025

- 11.05 750 10,580 n.d.

-9.95 800 33,262 n.d.

B. aannite = 0.1, asanidine -- 0 . 3 4 , amagnetite = 0.5 logfo~ -9.95

-8.9

-7.9

T (°C) fH~o ( b a r ) pH:O

800

850

900

1096 1350

3665 4050

11572 n.d.

i C a l c u l a t i o n m a d e o n t h e basis of e q u a t i o n 5 in t e x t (Wones a n d Eugster, 1 9 6 5 ) using f o - - T pairs f r o m curve 1 fig. 11.

34 PH20 values of around 5000 bar suggested b y the earlier calculations. What is more it is possible to achieve these values under conditions which are well supported' b y the mineralogical data. Enough evidence arises from the geothermometric calculations and other petrological arguments to suggest that the system is thermally heterogeneous (zoned) and therefore an exact T value is probably not very meaningful. These calculations certainly suggest that T is less than 900°C and that the activity of annite is significantly less than the simple mol. fraction of total iron. If for instance T = 850°C and the low value for the activity of annite (0.1) is accepted, then a pressure of nearly 5000 bar is calculated, in close agreement with the eruptive pressure calculated earlier. A number of models attempt to explain the occurrence of very violent volcanic eruptions. The simple rise of H20-bearing magma towards the surface with attendant saturation and boiling will n o t create overpressures of more than a few hundred bars. Models which involve mixing of separate batches of magma of different compositions, temperatures and water contents in the sub-volcanic conduits have recently received considerable attention (Eichelberger, 1978, 1980; Blake, 1981; Rice, 1981; Huppert et al., 1982). These may not be directly applicable to the Tambora case because: (a) they require active, largely open magmatic systems, or systems with short-lived closed chambers, whereas Tambora's lengthy period of inactivity suggests a sluggish, essentially closed system; (b) it is improbable that the mixing mechanism alone w i t h o u t substantial cooling and crystallisation will produce overpressures of several kilobars as required in this case. Numerous aspects of the petrology of the glassy trachyandesites erupted by Tambora in 1815 point to a closed system in which the water pressures built up gradually. These factors include: the occurrence of corroded, calcium-rich plagioclases, the inclusion of fluid/glass/vapour in clinopyroxene and the expansion of the field of biotite. More dynamic mechanisms such as magma mixing or the intrusion of sea water into shallow breached magma chambers are not favoured by these factors. The very nature of the eruption of Tambora suggests release of pressure from a closed system because of its very abrupt violent and short-lived character. The most obvious mechanism to explain such an eruption and the very high pressures involved is that of second boiling (Burnham, 1979a, b). This occurs when an H:O-bearing magma crystallises extensively and is unable to precipitate enough hydrated mineral species to prevent the increase in H20-content of the diminishing volume of residual melt. The process of second boiling is increasingly effective in generating high overpressures of water vapour at increasingly shallow depths (Burnham, 1979a, b). If crystallisation takes place in a closed magma chamber which is incapable of expansion then the total pressure is given by: 0.54 R T X m Pt =

V m (1 - X m)

(Burnham, 1979a, b)

(6)

35

Where Pt = total pressure, Vm = heat of fusion of crystalline phases (taken as 0.21 cal bar -1) and Xw m = mole fraction of H20 in the melt. Pt = Pin + P! where Pin is the internal overpressure and P1 is the lithostatic pressure. If T = 850°C, Pt = 5000 bars and Vm = 0.21 ca] bar -1, then X m is calculated as 0.466. Burnham (1975, 1979a, b) has developed a model for the dissolution of H20 in silicate melts based on the interaction of OH- molecules with the aluminosilicate melt structure. Thus the mole fraction of water which is soluble in a melt is determined relative to the mass of that melt (Me) equivalent to 100 g of NaA1Si3Os (Burnham, 1975, 1979a, b). The Me values for the major Tambora rock types are given in Table 13. The relationship of X~m to the weight fraction (Wm) of H20 dissolved in a silicate melt is given by (Burnham, 1975, 1979a, b) as: Xm

=

1- Zm

Me Ww m 18.02 (1 - Ww m)

for X m ~< 0.5

(7)

Figure 12 shows the relationship between X m and Ww m for the Tambora melts. The Me values for the range of compositions spanned by the Tambora lavas and glasses do not vary greatly and hence the curves in Fig. 12 are very similar for all Tambora rocks. The Xw m value of 0.466 calculated earlier can now be applied to eqn. (7) which then yields a value of 5.9 wt.% H20 in the 1815 magma immediately prior to eruption. Using this value it is then also possible to calculate the water content of the parent trachybasalt magma on the assumption that the enrichment of water during crystallisation of this to produce the trachyandesite was at least as much as that of other incompatible elements such as Rb and Zr. This factor averages about 1.8 and thus suggests that the trachybasalts had about 3.3 wt.% H20 (Xw m = 0.348). This can then roughly constrain the m i n i m u m emplacement level of the magma b o d y prior to the 1815 eruption to the lithostatic pressure (depth) for which the Xw m value of 0.348 yields an activity of H20 = 1. This is about 500 bar or approximately 1.5 km (see fig. 16.4, Burnham, 1979a). This is

/

°~

~ 04

o

~

~

TrachyandesiteT27

/ ~

J,

o,':

o'~o o~

o'~o o'~: o'o

Weight fraction H20 (Ww M)

Fig. 12. The variation of mole fraction of H20 (X m) in melts of trachybasalt and trachyandesite composition as a function of weight fraction (Wwm). Calculated using the model of Burnham (1975, 1979a, b).

36

because it is unlikely that the rising parent magma boiled initially as this would have lead to immediate eruption. If boiling did take place it would have continued until the water content of the liquid was reduced to that compatible with the appropriate lithostatic pressure (in this case 3.3% at 500 bar). The presence of chlorine in biotite and apatite and the presence of high C1 levels in the last crystallising liquid c o m p o n e n t of the intrusive rocks (indicated by the presence of sodalite) suggest that the liquids did not exsolve an H20-rich vapour phase prior to the isolation of the magmatic system to a closed chamber. This is because C1 partitions vary strongly to the hydrous vapour phase in favour of the silicate melt (distribution coefficients are of the order of 30 to 40 in favour of the H20-rich fluid) (Kilinc and Burnham, 1972). Thus the low C1 concentrations of the trachybasalt lavas indicate that these must have become vapour saturated at least at some point during their rise as liquids towards the surface. A M O D E L F O R T H E 1815 E R U P T I O N O F T A M B O R A

The caldera created by the 1815 eruption is a b o u t 6 km in diameter and 1 to 1.3 km deep. Barberi et al. (1983) indicate that the oldest lavas in the walls of the caldera are a b o u t 55,000 years old and that prior to 1815, the last major event was 5000 years earlier. The bulk of the 2000-m cone is composed of trachybasalt and low Mg trachybasalt lavas and intercalated pyroclastic rocks. According to the calculations presented earlier these rocks evolved by fractional crystallisation of magmas in an open system. They contained approximately 3% H20 and their temperatures were in the range 1100°C to 900°C. Fractional crystallisation produced evolutionary trends controlled by the separation of olivine, clinopyroxene, magnetite and plagioclase. The relative proportion of plagioclase in this assemblage was considerably lower than that involved in the generation of the trachyandesites later. The span of these broadly basaltic compositions generated in this stage of the petrogenetic cycle represent up to a b o u t 30% crystallisation. Of course, even the mos~ mafic trachybasalts sampled (PSS, Table 2) are relatively fractionated by comparison with likely primary basalts and must have suffered up to 20% crystallisation of olivine and clinopyroxene
37 basalt to low Mg trachybasalt. This magma could have been emplaced at depths as shallow as 1.5 km, in which case it would have been just water saturated, or at greater depths where it would have been initially undersaturated. The m a x i m u m depth of emplacement must have been a b o u t 4.5 km, based on the pressure required to saturate the trachyandesite liquid with 5.9 wt.% H20. The model then suggests that the magma b o d y further cooled as a closed system, crystallisation taking place under equilibrium conditions or a mixture of both equilibrium and fractional modes. The intrusive rocks (alkali gabbros or shonkinites and syenite) provide evidence first of crystaUisation and accumulation on the walls of the magma chamber, of phases equivalent to the phenocrysts in the trachybasalts (plagioclase, clinopyroxene, magnetite and olivine). These then reacted with increasingly hydrous and alkalirich residual liquids to produce biotite rims on the magnetite and olivine. By this stage ( a b o u t 70% crystallisation at the walls) the walls of the chamber will have reached the temperature of the trachyandesites as erupted (perhaps a b o u t 850°C) which probably increases into the interior of the body. If the magma was emplaced at the minimum depth (equivalent to 500 bar) a substantial excess water pressure will have been already generated. If it was emplaced at the m a x i m u m depth (equivalent to 1.5 kbar) water saturation will have just been reached in the wall regions where there has been extensive crystallisation. Further cooling of the magma chamber inwards from its walls proceeds until the outer wall temperature is a b o u t 700°C and this has resulted in the precipitation of interstitial alkali feldspar, sodic plagioclase and eutectic-style intergrowth of sodalite. At this stage the wall zone has solidified entirely and now forms a shell around the top and sides of the magma chamber. This situation is the same as that postulated by Burnham (1979b) to occur in high-level plutons supplying some p o r p h y r y copper systems. It is possible that this shell which lines the upper walls and t o p of the magma chamber provides both an impervious seal preventing gas leakage from the system as well as structural reinforcement of the chamber. At this stage when the walls of the chamber had reached 700°C the liquid in the interior had reached the composition of the glassy matrix of the trachyandesites (Table 11) and a temperature of a b o u t 850°C. This liquid contained 5.9 wt.% H20 and had evolved an aqueous fluid with a pressure of perhaps 4000 bar in excess of the lithostatic pressure. The eruption of T a m b o r a must have followed abrupt failure of the walls of the magma chamber and of the plugged conduit leading up to the summit of the volcano. By the time the eruption t o o k place the pressure within the chamber must have greatly exceeded the theoretical strength of the roof. It is possible that failure did n o t take place sooner because the quiet, closed magmatic system cooled slowly and only produced gradual increases in the rates of strain. H o w much expansion of the chamber t o o k place before erup-

38 tion is n o t known, but absence of any observed activity, rumbling or earthquakes until directly prior to eruption suggests that little or none occurred. The first real activity took place on 5 April and it must have been about this stage when the first fractures started to develop in the walls of the magma chamber. Once fracturing had started, then crack propagation would accelerate rapidly due to hydrofracture type phenomena. Due to the very high pressures involved and the enormous potential expansivity of the fluid phase, even the thinnest cracks reaching the surface would be the source of violent discharge such as that starting on 5 April. From this time on structural damage to the r o o f of the magma chamber would proliferate rapidly ending with the hypersonic discharge of the entire magma chamber which took place in the evening of 10 April and in the process disintegrated the upper 2000 m of the then volcanic cone. A Plinian column rose rapidly to great height and according to Stothers (1984) collapsed rapidly. ACKNOWLEDGMENTS The study was commenced while the author was the recipient of an Australian C o m m o n w e a l t h Government Postgraduate Award. In Indonesia the study was supported by the Indonesian Institute of Science (LIPI) with the cooperation of the University of Gadja Mada. The author thanks Dr. R. Varne of the University of Tasmania for his help and cooperation and Ms. B. Hosking, L. Lucas and S. Proferes at the University of Adelaide for their assistance in the production of the manuscript. Mr. J. Stanley and P. McDuie are also t h a n k e d for their contributions to the analytical work. Two anonymous reviewers are also thanked for their useful and constructive comments. REFERENCES Allard, P., Amy, B., Halbwachs, M., Matheron, S. and Vuillemin, J., 1983. The 1815 cataclysm of Tambora volcano, Indonesia: N e w evaluation of its volume, energy and acid gas output. Abstr. IAVCEI Symposium 18th IUGG, p. 38. Alzwar, M., Barberi, F., Bizouard, H., Boriani, A., Cavallini, A., Eva, C., Gelmini, R., Giorgetti, F., laccarino, S., Innocenti, F., MarineUi, G. and Sudradjat, A., 1981. A structural discontinuity with associated potassic volcanism in the Indonesian island arc: First results of the C N R - C N R S - V S 1 mission to the island of Sumbawa. Rend. Soc.

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