Magma mixing, fractional crystallization and volatile degassing during the 1883 eruption of Krakatau volcano, Indonesia

ELSEVIER

Journal of Volcanology

and Geothermal

Research 74 (1996) 243-274

Magma mixing, fractional crystallization and volatile degassing during the 1883 eruption of Krakatau volcano, Indonesia Charles W. Mandeville *, Steven Carey, Haraldur Sigurdsson University of Rhode Island, Graduate School of Oceanography, Narragansett, RI 02882-l 197, USA Received 4 December

1995; revised 30 July 1996; accepted 30 July 1996

Abstract The 1883 eruption of Krakatau volcano, Indonesia produced approximately 12.5 km3 (DRE) of magma, 90% rhyodacite, 4% mafic dacite, and 1% andesite (5% lithic material assumed). Melt compositions in the erupted products based on the study of glass inclusions and matrix glasses span a minimum range of 13 wt.% SiO,. Well defined linear trends of major elements have been successfully modelled by incremental fractional crystallization of plagioclase, pyroxene, and Fe-Ti oxides from andesitic parental magma to produce rhyodacite. This fractional crystallization model is supported by trace-element and isotopic data. The 1883 magma chamber was compositionally and thermally zoned with an upper portion of homogenous rhyodacite at a temperature of 880-890°C overlying more mafic dacite at 890-913”C, and andesite at 980-1000°C. Mafic dacite represents a hybrid magma formed by mixing small amounts of andesite with larger amounts of evolved rhyodacite. Evidence for magma mixing consists of banded pumices, glass inclusions that are more mafic than whole rock and matrix glass compositions, heterogeneous matrix glass compositions, and disequilibrium phenocrysts of anorthite-rich plagioclase in rhyodacite and mtic dacite. The 1883 rhyodacitic magma was probably not water saturated until it reached a shallow depth in the crust of 4-5 km. Glass inclusions indicate pre-eruption dissolved volatile content in rhyodacite and mafic dacite of 4.0 * 0.5 wt.%. Estimated sulfur discharge from erupted magma is 2.8 X 1012 g S, and estimated Cl discharge is 9.7 X 1012 g Cl. Two potential sources of additional sulfur from this eruption may be vaporization of seawater during entrance of pyroclastic flows into the sea, and degassing of nonerupted andesitic parental magma lying beneath evolved rhyodacitic magma in the zoned 1883 chamber. Keywords:

Krakatau; magma chambers; petrology: rhyodacite;

dacite; degassing;

1. Introduction The cataclysmic eruption of Krakatau volcano, Indonesia on August 26-27, 1883, is the second

* Corresponding author. Present address: Mineral Resources Department, Geological Survey of Japan, l-l-3 Higashi, Tsukuba, Ibaraki, 305 Japan. Tel.: 81-298-54-3634. Fax: 81-298-54-3533. E-mail: cmandy@?gsjrstn.gsj.go.jp.

volatiles; pbenocrysts;

P-T conditions

largest explosive eruption in historical times (Self and Rampino, 1981; Rampino and Self, 1982; Simkin and Fiske, 1983; Sigurdsson et al,, 1991a,b; Mandeville et al., 1996). It has been linked to a 0.3-0.4”C drop in mean annual northern hemisphere tropospheric temperature based on temperatqre records and dendrochronology studies in North America (Jones and Wigley, 1980; Rampino and Self, 1982; Papp, 1983; LaMarshe and Hirschboeck, 1984). Ex-

0377-0273/96/$15.00 Copyright 0 1996 Elsevier Science B.V. All rights reserved. PiI SO377-0273(96)000601

C. W. Mandeuille et al. /Journal

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of Volcanology and Geothermal Research 74 (1996) 243-274

plosive volcanic eruptions affect climate by altering the radiation budget of the Earth as a result of injection of large amounts of sulfur gases (mostly SO,) and silicate dust particles to stratospheric heights (Pollack et al., 1976; Toon and Pollack, 1982; Sigurdsson, 1990; Bluth et al., 1992). In order to evaluate the climatic impact of explosive volcanic eruptions it is thus necessary to quantify the source parameters for the event, which include the height to which tephra and gases are injected into the atmosphere, the mass of erupted tephra, and the total amounts of volatile components (principally SO, > that are released from erupted and non-erupted magma. At present, only preliminary estimates of these parameters exist for the 1883 Krakatau eruption. The only previous study of volatile discharge from the 1883 eruption was a reconnaissance study based on data from only two 1883 dacitic pumice samples (Devine et al., 1984). Despite its notoriety, no comprehensive petrologic study has been conducted on the entire range of compositions produced during the eruption. Previous studies by Camus et al. (19871, Self (19921, and a recent expedition to Krakatau (Sigurdsson et al., 1991a,b) have documented the occurrence of mixed gray and white pumices within both pyroclastic flow and fall deposits. These mixed andesitic to dacitic pumices indicate that a more mafic andesitic liquid may have been present in the 1883 magma chamber. In order to better define the total volatile yield, data from all magmatic compositions must be considered.

Table 1 Sample descriptions

In this paper, we present new bulk rock major and trace-element analyses of 1883 erupted products from a suite of samples that spans much of the eruption stratigraphy (Table 1; Sigurdsson et al., 1991a). In addition, we present new major and trace-element analyses of matrix glasses and glass inclusions from both evolved and more mafic magma types. We report new major-element analyses of mineral phases and use these data along with glass compositions to place constraints on pre-eruptive conditions within the 1883 magma chamber. We use the dissolved volatile concentrations found in glass inclusions and matrix glasses, together with more recent estimates of erupted volume (Mandeville et al., 1996), to estimate the minimum amounts of sulfur, chlorine and water degassing during the 1883 eruption.

2. Geologic setting Krakatau volcano is part of the Sunda Arc, that extends for 5600 km from the Andaman Islands in the northwest to the Banda Arc in the east. Subduction in this area is the result of convergence between the Indo-Australian plate and Southeast Asia. Krakatau lies within a volcanic belt trending N35”E within the Sunda Straits extending northward from the island of Panaitan to Krakatau, Sebesi, Sebuku, Radjabassa and the Sukudana plateau basalts in southern Sumatra (Fig. 1; Nishimura et al., 1986; Harjono et al., 1991). The Sunda Straits represents a

and locations

Sample a

Description

Location

KRA-129 KRA-141 KRA-142 KRA-045-l KRA-045-2 KRA-045-3 KRA-053 KRA-052 KRA-05 1 KIM-050 KRA-048 KRA-076

white rhyodacitic pumice from lithic and crystal poor pyroclastic flow deposit white rhyodacitic pumice from pyro- elastic flow deposit at beach level white rhyodacitic pumice from pyro- elastic flow deposit at beach level olive obsidian with pumiceous rind from basal pyroclastic flow same as above slightly vesicular olive obsidian gray dacitic pumice from basal pyroclastic flow deposit white rhyodacitic pumice from fall deposit white rhyodacitic pumice from fall deposit white rhyodacitic pumice from fall deposit bulk sample olive gray silty ash with accretionary lapilli and plant fragments 2 m wide pre-1883 andesitic dike

W Panjang west coast of Sertung west coast of Sertung SE Sertung Island same as above same as above W Rakata 500 m north of Island Point Island Point, W Rakata Island Point, W Rakata Island Point, W Rakata west coast Rakata Rakata cliff face

a Samples ordered according

to relative stratigraphic

position where possible (see Fig. 2).

C. W. Mandeville et al. /Journal

of Volcanology and Geothermal Research 74 (1996) 243-274

transitional area from oblique (55”) subduction under Sumatra, to near frontal (13” obliquity, Jarrard, 1986) subduction under Java (Newcomb and McCann, 1987; Diament et al., 1990; Harjono et al., 1991). Benioff zone depth beneath Krakatau is approximately 90 km (Newcomb and McCann, 1987). A microseismicity survey within the Sunda Straits (Harjono et al., 1991) reveals that crustal earthquakes in the Straits are clustered into three categories: (a) those located beneath the Krakatau volcanic complex and typically of tectonic origin; (b) inside a graben in the western portion of the Straits; (c) in a diffuse zone to the south of Sumatra. Focal mechanisms confirm that the Sunda Straits is under extension as a result of northwestward movement of the Sumatra fore-arc sliver plate along the Sumatra

245

Fault System (Curray et al., 1979; Huchon and LePichon, 1984; Lassal et al., 1989; Diament et al., 1990). The subsurface stratigraphy below Krakatau has been inferred from exploratory oil drilling to a depth of 3000 m, approximately 30 km southeast of the volcano (Fig. 1; Noujiam, 1976; De Neve and Sukardjono, 1985). The Pertamina/Aminoil C-l-SX well penetrated a marine succession of RecentPleistocene sediments, underlain by 2412 m of Upper Pliocene sedimentary rocks. This sequence is composed mostly of clays and claystones interbedded with quartz sandstone, siltstones and occasional tuffaceous beds. Granite and quartz monzonite xenoliths found within an 1883 pyroclastic flow unit on Sertung (Oba et al., 1982) and K-Ar dates of base-

107"E

5"3O'S-

Java Sea

Fig. 1. Location map of Sunda Strait region. Inset map shows present configuration of Krakatau islands and locations of samples (W ) listed in Table 1. Dashed box denotes trend of linear volcanic belt (N35oE). 0 denotes location of Pertamina/Aminoil C-l-SX well (see text).

246

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ofVolcanology and Geothetmal Research 74 (19961243-274

ment granitic rocks penetrated by oil wells off northwestern Java and southeastern Sumatra (Hamilton, 1979) indicate that the Sunda Straits region is most likely underlain by Cretaceous continental crust. Crustal thickness beneath Krakatau is approximately 22 km, as determined by the microseismicity survey of Harjono et al. (1991).

3. 1883 eruption

chronology

and stratigraphy

Eruptive activity at Krakatau began on May 20, 1883 with an explosive eruption from the Perboe-

watan vent, approximately 120 m above sea level (Verbeek, 1885). This eruption produced a minor pumice fall deposit approximately 30 cm thick near the vent, that was subsequently covered by 60 cm of dark gray fine ash (as described by Schuurman who visited the island on May 27 (Verbeek, 1885; Stehn, 1929; Simkin and Fiske, 1983). During the next three months, minor expiosive eruptions continued intermittently, producing minor amounts of ash fall on the Krakatau islands. By July, both Perboewatan and Danan vents were active. The island was visited by H.J.G. Ferzenaar on August 11 and he reported a 50-cm-thick layer of ash covering the beach, and his Panjang

Rakata 500 m South of island Point

East Sertung

KRA-052 -+

KRA-142

. KRA-141 KRA-045

pyroclastic surge

pyroclastic fall

KRA -048

1Ocm

lag breccia (not to scale)

phreatomagmatic ash fall

I Fig. 2. Stratigraphic columns from various locations on remnant Krakatau islands. Sample labels denote locations of samples measured sections, see Table 1 for sample descriptions. Pyroclastic flow deposits are not plotted according to 10 cm scale at left.

within

C. W. Mandeuille et al. / Journal of Volcanology and Geothermal Research 74 (1996) 243-274

map indicates at least three active vents (Verbeek, 1885; Sin&in and Fiske, 1983). The paroxysmal eruption began in the afternoon of August 26 when the intensity of explosive eruptions increased and an eruption column rose to at least 25-26 km based on the range of lithic fragments in pumice fall deposits and eyewitness reports (Simkin and Fiske, 1983; Sigurdsson et al., 1991a). The evolution of the 1883 eruption can be reconstructed from the succession of pyroclastic deposits on the Krakatau islands. Two widespread olive-gray and bluish-gray silty ash fall layers that overlie the 1883 soil horizon are the lowermost stratigraphic units of the 1883 deposits and are attributed to weak intermittent activity from May 20 to early August. These two fall deposits have been observed on the islands of Rakata (lo-25 cm thick), Panjang (5-10 cm thick) and Sertung (5-10 cm thick), and sometimes contain plant material, l-3 mm accretionary lapilli, and rare 3-4 mm dacitic pumice lapilli (Fig. 2; Mandeville et al., 1996). At some localities (e.g., west Rakata, Fig. 2) the layers have a slightly vesiculated texture. Onset of the catastrophic August 26-27 eruption is represented by deposition of a coarse-grained, light gray Plinian pumice fall unit, 5-20 cm thick, which overlies the olive-gray silty ash fall layers on Rakata, Sertung and Panjang. Above this pumice fall deposit are interbedded surge and fall deposits, which are generally l-3 m in total thickness, but locally up to 4-6 m in thickness. Occasional small, < l-mthick, pyroclastic flow units are interbedded with the surge and fall deposits and most likely represent short-lived instabilities in the eruption column rather than wholesale column collapse, as evidenced by overlying 0.2-0.7-m-thick pumice fall units on Rakata, Panjang and Sertung. Transition to sustained column collapse and pyroclastic flow generation is probably contemporaneous with the first of several large explosions in the early morning on August 27, and indicated by the 40-60-m-thick accumulations of multiple pyroclastic flow units in the upper portions of the stratigraphy on all proximal Krakatau islands (Self and Rampino, 1981; Sigurdsson et al., 1991a). Emplacement of a widespread, voluminous high energy flow most likely was coincident with the largest explosion which occurred at 10:00 am on August 27 and may have been related to an increase

247

in magma discharge rate as a result of downward collapse of the main portion of Krakatau island to form a submarine caldera (Carey et al., 1996).

4. Description of 1883 erupted products Highly vesiculated white pumice constitutes over 85% of the 1883 erupted products. This pumice contains 7-15 vol.% phenocrysts (vesicle-free basis), which in order of decreasing abundance consist of euhedral-subhedral plagioclase, orthopyroxene, titanomagnetite, clinopyroxene and ilmenite (Table 2). Titanomagnetite and ilmenite phenocrysts also occur as inclusions within plagioclase, orthopyroxene, and clinopyroxene. No exsolution lamellae were observed in the Fe-Ti oxide phases and titanomagnetite is approximately three to four times more abundant than ilmenite. Acicular apatite crystals and pyrrhotite blebs are common accessory phases occurring mostly as inclusions within phenocryst phases. Matrix glass in the white pumices is clear, and highly vesiculated. Phenocrysts in the white pumices are commonly present as 2-3-mm clots giving the pumice a glomeroporphyritic texture. Rare isolated blebs of gray glass containing abundant titanomagnetite microlites f phenocrysts also occur in some white pumices. Olive-colored obsidians (Tables 1 and 2) comprising from 1 to 8% of some 1883 pyroclastic deposits (Mandeville et al., 1996) contain 13-15 vol.% pheTable 2 Modal analyses of Krakatau

pumice and obsidian

samples

Sample

Plag.

Opx.

Cpx

Tmt.

Ilm.

Glass

Other a

KRA-045-l KRA-045-2 KRA-045-3 KRA-050 KRA-05 1 KRA-052 KRA-053-2A KRA-053-l KRA-129 KRA-141 KRA-142 KRA-076

9.7 8.4 8.7 6.6 1.9 9.0 8.2 2.1 1.6 4.6 3.4 18.3

1.7 3.3 1.1 tr. 0.2 1.4 2.2 0.5 0.7 0.5 1.0 4.4

0.6 1.5 0.8 0.3 0.5 tr. 0.4 0.3 0.1 0.1 1.6 1.1

1.1 1.6 0.9 0.3 0.9 1.0 1.0 1.8 0.2 0.3 0.5 2.5

tr. tr. tr. tr. tr. tr. tr. tr tr. 0.1 0.1 0.8

86.7 85.0 88.5 92.0 96.5 88.6 85.1 94.3 97.3 94.4 93.4 72.8 b

0.2 3.1 1.0 0.1

a Denotes content of matic xenolithic inclusions, apatite. b Microlitic groundmass containing plagioclase, clinopyroxene, titanomagnetite and ilmenite.

pyrrhotite

and

orthopyroxene,

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C. W. Mande~~ille et al. / Journnl cf Volcanology und Geothetmul

nocrysts (vesicle free) in order of abundance of plagioclase, orthopyroxene, titanomagnetite, clinopyroxene, ilmenite, apatite and pyrrhotite. They also contain phenocryst clots similar to those observed in the white pumices. Matrix glass in the olive obsidians is clear where slightly vesiculated, to slightly yellow where dense and vesicle poor. Well aligned microphenocrysts of plagioclase and clinopyroxene define a flow fabric that is common in these obsidians. Some obsidians exhibit a gradual transition in vesicularity and color into white pumice on the scale of 10 cm. Light to dark gray pumices are also present in both pyroclastic fall deposits and flow deposits (Table 1). These gray highly vesiculated pumices comprise about 4% of the 1883 erupted products and contain 4- 12% phenocrysts of plagioclase, orthopyroxene, titanomagnetite, clinopyroxene, ilmenite, and irregular cognate(?) lithic fragments and/or chilled mafic inclusions. Matrix glass in the gray pumices is colorless, but contains abundant microlites of titanomagnetite which gives the glass its gray color on macroscopic examination (Verbeek, 1885). These titanomagnetite microlites give the gray pumice higher magnetic susceptibility than white pumice of equal volume (Mandeville et al., 1994). Phenocryst clots present in the gray pumices produce a slightly glomeroporhyritic texture. The color of the matrix glass in hand specimens of gray pumice is heterogeneous and can range from light to very dark gray. In addition, isolated 2-3-mm-size patches of clear matrix glass can be found with sharp to gradational contacts within gray pumice hosts. Strongly banded pumices comprise less than 1% of 1883 erupted products, and these contain varying proportions of clear and gray glass. Contacts between clear and gray matrix glass can be sharp in some samples indicating mingling of different magmas to completely gradational in others, indicating mixing of magmas to form hybrid compositions (Anderson, 1976; Hildreth, 198 1; Bacon, 1986; Sparks and Marshall, 1986).

5. Samples and procedures Samples chosen for this study span a significant portion of the 1883 eruption stratigraphy. They in-

Research 74 (19961 243-274

elude a basal olive-gray silty ash layer attributed to low level phreatomagmatic activity from May 20 to early August (Sigurdsson et al., 199la), samples from three successive pumice fall deposits (attributed to plinian activity on August 26-271, and obsidian and pumice samples from the main pyroclastic flow phase of the eruption (see Figs. 1 and 2 and Table 1 for description and location). Modal counts of 1000 points were done on one to three polished thin sections from each sample and the proportions of glass and mineral phases were calculated on a void-free basis (Table 2). Polished grain mounts of plagioclase, orthopyroxene, clinopyroxene, and Fe-Ti oxides from each sample were prepared for microprobe analyses. Standard polished thin sections were also used for matrix glass, glass inclusion and mineral analyses by electron microprobe. Rock powders for XRF determinations of major and trace elements were prepared by crushing cleaned samples in a steel mortar then grinding and homogenizing in an agate mortar. Major- and trace-element analyses of whole-rock samples by XRF were conducted at the University of Rhode Island according to Danforth (1986) and Hamidzada (1988). U.S. Geological Survey rock standards AGV-1 and W-2 were analyzed in duplicate as monitors for precision and accuracy. Major-element analyses of matrix glasses, glass inclusions and minerals were conducted at Brown University using a Cameca Camebax electron microprobe. An accelerating voltage of 15 keV, 10 nA beam current, and 10 s count time were used for glass analyses. Beam diameter was set at approximately 10 pm in order to minimize Na loss (Nielsen and Sigurdsson, 1981). Precision and accuracy were monitored by repeated analyses of RN-18, a comenditic obsidian standard. Major-element analyses of minerals were obtained using an accelerating voltage of 15 keV, 15 nA beam current, 10 s count time, and focused beam. Natural mineral and glass reference standards were used in the calibration procedure and to monitor for instrumental drift. Sulfur and chlorine analyses in matrix glasses and glass inclusions were obtained utilizing a 15 keV accelerating voltage, 30 nA beam current, 5-10 pm beam diameter, and 50- 100 s count time. Peak searches were conducted on three separate plagio-

C. W. Mandeuille et al. / Journal of Volcanology and Geothermal Research 74 (19%) 243-274

clase-hosted glass inclusions in samples KRA-053C and KRA-053A, in order to determine the wavelength of the SKcr peak (S*- or S6+, Carroll and Rutherford, 1988). N.I.S.T. Standard Reference Material 620 (S-bearing soda-lime glass, 1155 ppm S, present as S6+ only) and troilite (FeS) were used as calibration standards. Prior to analysis of unknowns, the sin 8 address was reset in accordance with the S peak searches on the unknowns. Scapolite reference standard ON70 (1.9 wt.% Cl, Evans et al., 1969) was used to standardize Cl. Whole-rock and glass-separate trace-element analyses were conducted by XRF at the University of Rhode Island, by ICP-MS at Harvard University and the Geological Survey of Japan, and by ICP-ES at Boston University. Trace-element analyses on glass separates by ICP-ES at Boston University were done according to Murray et al. (1995). Sample prepararation for ICP-MS REE analyses at Harvard was done according to Murray and Leinen (1993). Replicate analyses (n = 3) of U.S. Geological Survey rock standard AGV-1 (andesite, Gladney et al., 1983), and an internal reference sample BT-15 (Bishop Tuff, Gardner et al., 1991) were used as monitors for precision and accuracy of both ICP-MS and ICP-ES analyses. ICP-MS analyses done at the Geological Survey of Japan followed the methods of Imai (1990). Replicate analyses (n = 3) of Geological Survey of Japan Geochemical Reference Samples basalt JB-2, and andesite JA-1 (Imai et al., 1995), were used as monitors for precision and accuracy.

6. Major-element chemistry of whole-rock samples

these as rhyodacitic pumices. Al,O, content is slightly variable ranging from 14.7 to 15.3 and probably results from variable crystal content (Table 3; Fig. 3). Total alkali content (Na,O + K,O) ranges from 6.87 to 7.24. FeO* content of white rhyodacitic pumices ranges from 3.13 to 3.73 Whole-rock chemical data for olive obsidian samples are similar to that of the white rhyodacitic pumices, particularly to samples that have at least 10% crystallinity (e.g., KRA-052). Major-element compositions of obsidian samples lie within the composition range of rhyodacitic pumices and one sample extends to slightly lower SiO, content (Table 3; Fig. 3). Weight percent SiO, in obsidians ranges

0 whRe n

rhyodacite

olive obsidian

A gray da&e

67

66

69

70

71

Weight % S102

~JF-yy-l

E

m

.

Q 2.7

3

0

2’567

I 71

3.0 001 2.0 ; 2.6 z 2.4 Q 2.2 3

Major-element compositions of the 1883 Krakatau products are listed in Table 3 along with published XRF analyses from Self (19921, and two analyses from Stehn (1929). We have assigned rock names based on recalculated (anhydrous) SiO,: basalts contain < 52% SiO,; basaltic andesites, 52-57% SiO,; andesites, 57-64% SiO,; dacites 64-69% SiO,; rhyodacites 69-72% SiO,; rhyolite, > 72% SiO, modified after LeBas et al. (1986). The range in SiO, of white 1883 pumices is relatively small from 68.8 to 70.1% spanning the dacite to rhyolite field boundary of LeBas et al. (1986). We will subsequently refer to

249

2.0 1.6 67

144 67

66

’

68

69

*

70

69 70 Weight % SIO,

71

’

71

Fig. 3. Major-element variation diagrams for (a) Fe0 versus SiO, , (b) CaO versus SiO,, (cl K,O versus SiO, and (d) Al,O, versus SiO,, all data recalculated to volatile-free oxide sums of 100%.

250

C. W. Mandeuillr et al./ Journal of Volcanology and Geothermal Research 74 (1996) 243-274

oxides, though extending the range to slightly SiO, and higher FeO’ and MgO contents. variations in major-element chemistry in gray pumices also could be attributed to variable and mafic inclusion contents (Table 2).

from 68.3 to 69.7. The olive obsidians are the quenched equivalent of the white rhyodacitic pumices. Samples of gray 1883 pumice are slightly more mafic than white rhyodacite, with SiO, varying from 68.1 to 68.9% and we will subsequently refer to these as gray dacitic pumices (Table 3; Fig. 3). We have not analyzed gray pumice sample KRA-053C for major elements, though it would be more mafic than the others based on its matrix glass chemistry (Table 4) and crystal content. Two gray pumice whole-rock analyses (#‘s 2083 and 2317, Table 3) from Stehn (1929), when recalculated on an anhydrous basis have 67.7-69.2 wt.% SiO,. These analyses of Stehn (1929) are within the range of most

Table 3 Major- and trace-element Sample: Code d:

composition

045-2

045-3

00

of whole-rock

lower Slight dacitic crystal

7. Mineral chemistry Most plagioclase phenocrysts (> 75%) in white rhyodacitic pumice samples are normally zoned with average rim composition of An,, and core compositions of An,,_,,. Unzoned to reversely zoned plagioclase phenocrysts are common but less abundant. Representative plagioclase compositions for white

samples

052 wrd

053D

053A gd

141 wrd

gd

7’ 00

076 a

2083 8

gd

129 v+rd

2317 g

00

050 wrd

5s

gd

gd

68.98 0.76 14.99 3.4s 0.14 0.8 I 2.84 4.73 2.21 0.16 99.04

66.86 0.75 14.54 3.42 0.14 I.10 3.02 4.92 2.26 0.16 97.17

68.3 I 0.76 14.94 3.33 0.14 0.93 3.07 4.9 I 2.24 0.17 98.78

67.74 0.76 15.30 3.67 0.14 I .03 3.04 5.47 2.15 0.17 99.44

68.21 0.78 IS.18 3.86 0.15 I .09 2.97 5.17 2.0’) 0.17 Y9.64

69.60 0.74 14.70 3.24 0.13 0.83 2.58 4.53 2.33 0.16 98.82

6X.41 0.79 15.17 3.70 0.14 0.92 3.03 4.53 2.29 0.16 99.13

67.5 I 0.71 14.41 3.61 0.14 I.14 2.65 5.26 2.59 0.10 98.15

67.43 0.75 15.54 3.51 0.15 0.76 2.92 5.16 2.38 0.09 98.69

64.08 0.88 16.58 5.38 0.14 1.51 4.45 4.01 1.90 0.35 99.27

67.64 0.57 14.54 3.92 0.06 0.99 3.02 4.03 2.91 0.06 97.74

66.38 0.64 16.94 5.21 0.08 I.14 3.11 4.09 I .78 0.00 99.37

31 358 I80 34 58

4Y 358 200 32 61 6 251 47 b.d. 76 b.d.

36 366 I83 32 63 b.d. 234 38 20 83 I9

3x 364 I95 38 66 6 246 50 IO 134 b.d.

61 392 I84 35 63

57 366 204 32 62 8 252 54 b.d. Y6 b.d.

_

_ _ _ _ _ _ _ _ _ _ _

Oxide (wt. 5%)a TiOt SiO At&& Fe0 * MnO MgO CaO Na,O RzO PZO, Total

69.23 0.73 15.07 3.37 0.14 0.8 I 2.86 4.74 2.31 0.16 99.40

Trace elements fmnc. Rb 61 Ba 380 Sr 205 La 42 Ce 69 Nb 8 Zr 268 Y 56 b.d. CU Zn 70 Ni b.d.

ppml ’ 59 371 202 38 64 x 269 56 b.d. 68 b.d.

228 33 b.d. 62 b.d.

268 56

b.d.

_

_

_

-

-

44 345 318 31 62 4 I98 43 82 19

Fe0 * = Total Fe as FeO: b.d. = below detection (e.g., for Cu < 2 ppm; for Ni < 10 ppm). a Major- and trace-element chemistry done by XRF at University of Rhode Island. ’ Major-element XRF analysis of 1883 gray pumice from Self (1992) recalculated. ’ Major-element XRF analysis of 1883 obsidian from Self (1992) recalculated. d Sample code: oo = olive obsidian; wrd = white rhyodacite; gd = gray dacite; a = pre-1883 andesite from andesite dike. ’ Trace-element XRF analyses done at University of Rhode Island. ’ Gray dacitic pumice analyses of Stehn (1929) recalculated.

C. W. Mandeuille et al. / Journal of Volcanology and Geothermal Research 74 (19%) 243-274

251

Table 4 Major- and trace-element composition of matrix glasses and glass inclusions Sample: code’:

045-l mg ’ 00

Oxide Cwt.%) d TiOi SiO 70.83 0.65 (09) (1.2) e 4203

Fe0 * MnO MgO CaO Na,O K2O

Total VBD n

f

14.46 2.70 0.14 0.70 2.06 5.46 2.37 99.37 0.63 6

(32) (28) (02) (10) (27) (16) (10) (32)

045-2 mg 00

045-3 mg oo

045gib 00

048 gi wrd

048 mg wrd

048 gi a

050 mg wrd

050 gi wrd

72.17 0.53 (02) (22)

71.33 0.55 (41) (07)

68.87 0.48 (66) (03)

69.04 0.51 (10) (65)

71.54 0.39 (34) (04)

60.43 1.26 (02) (07)

72.82 0.53 (45) (10)

69.00 0.57 (16) (86)

14.07 (08) 2.62 (07) 0.13 (05) 0.57 (03) 1.79 (06) 5.27 (28) 2.46 (13) 99.62 (09) 0.38 3

14.14 (15) 2.41 (10) 0.09 (05) 0.59 (03) 2.10 (08) 5.29 (16) 2.42 (08) 98.93 (51) 1.07 15

13.17 (29) 2.31 (13) 0.13 (04) 0.56 (02) 1.68 (06) 5.03 (44) 2.37 (11) 94.59 (1.1) 5.41 4

13.40 (45) 2.14 (36) 0.03 (03) 0.54 (12) 1.84 (28) 4.99 (24) 2.43 (10) 94.92 (88) 5.08 3

13.25 (10) 2.19 (12) 0.08 (11) 0.47 (05) 1.74 (09) 5.21 (16) 2.52 (05) 97.38 (18) 2.62 3

14.55 (24) 6.75 (52) 0.16 (01) 1.88 (05) 4.44 (88) 4.84 (32) 1.94 (09) 96.25 (44) 3.75 3

13.79 2.32 0.09 0.57 2.06 5.54 2.29

13.60 2.64 0.12 0.60 2.19 4.82 2.37

Trace elements fconc. ppm) g Rb 64 Ba 403 Sr _ 175 La _ 42 Ce _ 73 Nb 9 Zr 284 Y 62 _ CU b.d. Zn 72 Crh 3.4 Nib 12.8 _ co h 2.6

68 396 178 38 64 9 292 62 b.d. 70 -

_ _ _

_ _ _ _ _ _ _

_

rhyodacite samples are presented in Table 5. A strong mode in both rim and core compositions (Fig. 4a) indicates the relatively uniform nature of the plagioclase phenocryst population. Minor tails extending to more albite-rich rims may be the result of plagioclase growth during final stages of decompression. The presence of plagioclase phenocryst rim compositions extending to An,,_,O, and core compo.. smons to An,_,, may reflect inheritance of plagioclase phenocrysts from an andesitic parent magma, as these compositions would be out of equilibrium with rhyodacitic to rhyolitic melt at low water pressures near 1.O- 1.5 kbar (Gill, 1981; Rutherford et al., 1985; Rutherford and Devine, 1988; Housh and Luhr, 1991; Nakada et al., 1994). Plagioclase phenocrysts in gray dacitic pumices are both reversely zoned and normally zoned in approximately equal proportions, and unzoned phe-

_ _ _ _ _ _

_ _ _

_ _ _ _ _ _

(15) (27) (03) (03) (10) (13) (03)

lCW0 (04) 0.00 3

61 403 169 33 61 8 275 60 56 90 3.3 14 2.3

(18) (20) (04) (05) (24) (07) (10)

95.91 (30) 4.09 5

-

nocrysts are common but less abundant. Representative normally zoned plagioclase is presented in Table 5, and a histogram of plagioclase rim and core compositions is shown in Fig. 4b. Plagioclase phenocrysts in the gray dacite exhibit strong modes in rim and core compositions between 46 to 50% An (Fig. 4b). An average rim composition for a composite of four gray dacite samples is An,,,, (1 o = 5.2, n = 50). Two gray dacite samples have average plagioclase rim compositions of An,,,O_,,.,, higher in An content than averages observed in white rhyodacitic samples. Reversely zoned plagioclase phenocrysts are more abundant in gray dacite samples. Unlike the plagioclase from the white rhyodacitic samples, there is a greater abundance of plagioclase rim and core compositions of An56_65,in gray dacitic pumice. Rare An,,_,, rim and core compositions in gray dacite are probably xenocrysts inherited from

C. W. Mandeville et al. / Joumul of’ Volcanology and Geothermal Research

252

74 (1996) 243-274

Table 4 (continued) Sample: Code ‘:

051 mg a wrd

Oxide (wt. %) * SiO, 7 I .97 TiO? 0.42 Al,& 13.59 2.41 Fe0 * MnO 0.11 MgO 0.56 CaO 1.88 Na,O 5.22 KzO 2.42 Total 98.57 VBD ’ 1.43 n 12

(58) ’ (06) (26) (19) (04) (05) (16) (30) (10) (32)

051 gi h wrd

69.50 0.54 13.21 2.27 0.10 0.62 1.93 5.03 2.41 95.60 4.40 10

Truce r1ement.s fconc. ppmi g Rb Ba _ Sr _ _ Nb _ _ Zr _ _ Y CU Zn Cr h _ _ Ni h co h

(1.2) (12) (40) (25) (04) (14) (17) (34) (05) (09)

052 mg w*rd

72.04 0.56 13.65 2.49 0.09 0.58 2.16 5.27 2.53 99.37 0.63

(59) (04) (17) (14) (03) (02) (17) (20)

(I?) (66)

052 gi wrd

69.64 (96) 0.49 (14) 13.46 (30) 2.27 (36) 0.12 (04) 0.57 (14) 1.99 (20) 4.85 (31) 2.50 (21) 95.89(1.1) 4.1 I

0

63 405 175 279 58 IX 74 3.6 12.9 ,?

more mafic parent magmas or cumulate material commonly found in zoned talc-alkaline magmatic 1976; Gill, 1981; Druitt and systems (Anderson, Bacon, 1989; Sigurdsson et al., 1990a; Nakada et al., 1994). Most orthopyroxene phenocrysts in white rhyodacitic pumices show slight reverse zoning, with rim composition Mg#‘s exceeding those of cores by 1.0-2.2. One orthopyroxene rim and core pair exhibits normal zonation, and one pair is unzoned. Average Mg# of rims is 70.7 + 0.6, whereas average core Mg# is 69.6 f 0.8 (Table 6). All orthopyroxene phenocrysts in the gray dacite show slight reverse zoning, with average Mg# of rims and cores of 71.1 + 0.6 and 70.2 + 0.8, respectively. These average values are slightly higher than was observed in white rhyodacite samples (Table 6). Orthopyroxene phenocrysts in gray dacite contain higher TiO, and MgO than those in rhyodacite, and slightly lower Fe0 (Table 6). Five out of nine clinopyroxene rim and core pairs

053A mg

053A gi

053.2A mg

053-2A gi

053C mg

gd

gd

gd

gd

gd

69.69 0.58 13.45 2.92 0.21 0.81 2.05 3.90 2.49 97.10 2.90 II

(2.4) (08) (86) (85) (13) (67) (69) (59) (47) (75)

69.17 (1.3) 0.56 (14) 13.21 (1.7) 3.33 (58) 0.13 (05) 0.58 (18) I .70 (66) 4.97 (82) 2.52 (38) 95.17(1.1) 4.83

69.72 (2.5) 0.49 (12) 14.80 (1.7) 2.23 (40) 0.12 (01) 0.44 (23) 4.29 (1.6) 5.20 (74) 2.60 (70) 99.89 (I.11 0.1 1 4

h.5

_

3x3 IXI 0 270 57 63 I I7 (3.6 Il.7 3.3

_

_ _

69.76 0.53 12.45 2.44 0.13 0.65

(1.2) (08) (36) (23) (06) (10) 1.48(28) 4.65 (21) 2.56 (13) 94.66 (04) 5.34 9

_

_ _ _ _ _ _ _ _

68.74 (1.6) 0.69 (10) 14.50 (37) 3.19 (51) 0.15 (03) 1.03 (23) 2.98 (36) 5.38 (30) 2.14(11) 98.78 (66) 1.22 5

_

_ _ _ _ _ _ _ _ _

white rhyodacite are slightly reversely zoned, three pans are normally zoned, and one pair is unzoned. Average Mg# of clinopryoxene rims is 77.6 and Mg# of cores is 77.1 (Table 6). Clinopyroxene phenocrysts in gray dacite show slight normal zoning, and less commonly weak reverse zoning. Four phenocrysts are normally zoned and three are reversely zoned. Average rim Mg# is 76.1 + 1.0, and average core Mg# is 76.1 + 1.5 (Table 6). TiO, contents of clinopyroxene in gray dacite are higher than in rhyodacite (Table 6). Compositions of titanomagnetite and ilmenite phenocrysts in white rhyodacite are presented in Table 7. Average ulvospinel content in titanomagnetite is 26.3 mole % (l(+= 0.8, n = 28) and the range is from 24.2 to 27.0 when calculated according to the method of Stormer (1983). Average mole% ilmenite in the rhombohedral phase is 74.2 (1 cr = 0.8, n = 25) with a range of 73.2 to 77.1. Titanomagnetite compositions in the gray dacite are slightly more titanium rich than in white rhyodacitic samples iI1

C. W. Mandeville et al. / Journal of Volcanology and Geothermal Research 74 (1996) 243-274

253

Table 4 (continued) Sample: Code ‘:

403

Fe0 * MnO MgO CaO Na,O I’+ Total VBD f n

141 mg wrd

141 gi wrd

141 mg a

141 gi a

142 mg wrd

142 gi wrd

72.78 (67) 69.07 (1.31 0.56 (04) 0.61 (15) 13.88 (32) 13.74 (26) 2.29 (14) 2.57 (47) 0.12 (02) 0.10 (05) 0.59 (03) 0.67 (121 1.97 (121 2.10 (30) 5.50 (37) 5.10 (18) 2.35 (12) 2.36 (11) 100.04 (92) 96.33 (46) -0.04 3.67 10 9

69.83 (1.0) 0.52 (10) 13.93 (28) 2.51 (24) 0.13 (03) 0.70 (16) 2.28 (24) 5.41 (34) 2.36 (16) 97.65 ( 1.2) 2.35 4

67.16 (1.2) 0.64 (16) 13.62 (37) 2.74 (31) 0.11 (041 0.76 (14) 2.36 (26) 4.95 (28) 2.28 (10) 94.62 (73) 5.38 12

63.37 (94) 1.04 (11) 15.30 (10) 5.28 (72) 0.17 (02) 1.98 (39) 4.55 (47) 5.46 (09) 1.52 (27) 98.67 (63) 1.33 3

60.48 (1.8) 1.14 (16) 14.42 (66) 5.76 (1.1) 0.15 (05) 2.26 (58) 4.30 (37) 4.97 (58) 1.50 (07) 94.97 (1.1) 5.03 5

71.71 (56) 0.55 (05) 13.93 (24) 2.35 (13) 0.10 (08) 0.58 (03) 1.74 (06) 5.38 (27) 2.39 (08) 98.74 (79) 1.26 12

69.57 (1.3) 0.52 (12) 13.51 (32) 2.25 (37) 0.07 (03) 0.55 (15) 1.66 (26) 4.98 (32) 2.41 (05) 95.52 (74) 4.48 7

129 mg a wrd

Oxide (wt.%) d 67.97 (1.6) ’ SiO,

TiO,

129 gi b wrd

053C gi gd

0.66 (28) 13.79 (35) 2.60 (40) 0.13 (05) 0.74 (16) 2.33 (34) 5.21 (43) 2.25 (15) 95.67 (1.1) 4.33 11

Trace elements (cont. ppm) g

Rb Ba Sr La Ce Nb Zr Y CU Zn Crh Nib Cob

_ _ _ _ _ _ _ -

62 394 170 39 70 9 274 59 53 76 3.7 11.7 2.4

_ _

64

_ _ _ _ _ _

_ -

394 174

_ _

34 64 8

-

276 58 b.d. 88 14.1 18.1 6.6

_ _ -

-

Fe0 * = Total Fe as FeO; n = number of electron microprobe analyses; b.d. = below detection (e.g., for Cu < 2 ppm); a mg denotes matrix glass. b gi denotes glass inclusion. ’ Sample code: oo = olive obsidian; wrd = white rhyodacite; gd = gray dacite; a = andesite. d Major-element chemistry done by electron microprobe, Brown University. e Numbers in parentheses express 1o of the data in smallest units of value at left. f VBD = volatiles by difference (100.00 - analytical total). g Trace elements by XRF at University of Rhode Island. h Trace elements by ICP-MS Harvard University (R. Murray, analyst).

(mole% ulvospinel from 26.4 to 29.6, Table 7). Mole% ilmenite in gray dacitic samples is 74.3 to 74.4 (Table 7).

8. Major-element

chemistry of 1883 matrix glasses

Matrix glass compositions within the white rhyodacitic pumices are rhyodacitic-rhyolitic with SiO, (anhydrous) ranging from 70 to 74% (Table 4; Fig. 5). Oxide variation diagrams plotted versus SiO, for

FeO’, MgO, TiO,, CaO and K,O exhibit well defined linear trends (Fig. 5). Our range of values for Na,O are higher than those reported for Krakatau 1883 matrix glasses by Camus et al. (1987) but are within the range reported by Devine et al. (1984) because we have corrected for Na loss during electron bombardment (Nielsen and Sigurdsson, 1981). There is a slight decrease in Na,O with increasing SiO, and a moderate decrease in Al,O, with increasing SiO,. One white rhyodacite pumice sample, KRA-141 contains colorless andesitic glass adhering to plagio-

254

C. W. Mandeoille et al. /Journal

of Volcanology and Geothermal Research 74 C19961243-274

clase phenocrysts (63% SiO,; Table 4; Fig. 5), and white rhyodacitic pumice sample KRA-052 contains isolated patches of gray dacite. Matrix glasses in the three samples of olive obsidian (Table 4) are relatively homogenous and similar in composition (for all oxides) to glasses in rhyodacitic pumices (Fig. 5). Matrix glass adhering to plagioclase phenocrysts hand picked from the olive gray silty ash layer (KRA-048, Table 4) is similar in composition to matrix glasses from white rhyodacite and obsidian. Matrix glasses from gray pumices are generally more variable in SiO,, Al,O,, CaO, Fe0 * , MgO and TiO, content on the scale of a thin section than most white rhyodacitic pumices. SiO, ranges from 67.3-74.5 with an average value of 70.8 (1 (T = 2.47. n = 20). However, part of the increased range to evolved SiOz content of 74.5% originates from normalization of gray matrix glass analyses with analytical totals below 98 wt.% to 100%. Although the gray dacite matrix glasses are similar in terms of their average oxide values relative to white rhyodacitic matrix glasses (except sample KRA-053C. Table 4), gray dacite matrix glasses show larger ranges in FeO*, MgO, CaO, K,O, A120,. TiO,, and higher MnO content than rhyodacitic matrix glasses (Fig. 5). Avoidance of titanomagnetite microlites in the gray matrix glass during probe analyses

Table S Representative

plagioclase

Sample: KRA-050 Code: wrd Rim Oxide SO,

KRA-053C gd

Rim

Rim

Core

Core

Rim

Core

Cwt. %i ”

K# Total

58.29 28.83 0.26 9.28 5.54 0.18 102.38

XA” XAb X or

0.48 0.51 0.01

AW, FeO’ CaO Na,O

Core

KRA-053A gd

55.50 56.32 54.69 55.18 55.23 56.72 53.72 27.16 27.41 27.95 27.73 29.15 27.99 29.89 0.48 0.44 0.42 0.54 0.47 0.52 0.52 10.11 9.56 10.86 10.02 11.24 9.88 11.67 5.35 5.65 5.30 5.23 4.76 5.55 4.56 0.17 0.20 0.13 0.21 0.16 0.18 0.12 99.37 99.58 99.35 98.91 101.01 100.84 100.48 0.51 0.48 0.01

0.48 0.51 0.01

0.54 0.45 0.01

0.51 0.48 0.01

0.56 0.43 0.01

0.49 0.50 0.01

2

25:

s =I

20:

$ t

15:

Mole % Anorthite 35

qgray core H gray rims

30

10 5 0

Mole % Anorthite Fig. 3. (a) Plagioclase composition histogram from white rhyodacitic samples KRA-050, KRA-052. KRA-129, KRA-142. (b) Plagioclase composition histogram from gray dacitic samples KRA-053A, KRA-053C, KRA-053A-IC, KRA-053.2A; Representative analyses shown in Table 5.

cornposItions

KRA-052 wrd

qwhite cores

0 white rims

0.58 0.41 0.01

Fe0 * = total Fe as FeO; X,, = mole fraction anorthite: X,, = mole fraction albite; X0, = mole fraction orthoclase. a Plagioclase analyses by electron microprobe, Brown University.

has most likely lead to an underestimation of Fe0 * , TiOz, and MgO contents. Bulk chemical analyses of gray matrix glasses would probably show even greater compositional differences in terms of Fe0 * , TiO, and MgO content.

9. Composition

of glass inclusions

Glass inclusions are present in plagioclase, clinopyroxene, orthopyroxene, and titanomagnetite phenocrysts in all products of the 1883 eruption. They range from 10 to 150 pm in diameter and vary in color from dark olive brown to pale gray, or

C. W. Mandeville et al. /Journal

255

of Volcanology and Geothermal Research 74 (19%) 243-274

colorless. These glass inclusions are generally not devitrified, and contain one or more vapor contraction bubbles, and occasional euhedral microlites of apatite, Fe-Ti oxides and pyrrhotite, similar to those described in glass inclusions by Anderson (1974) and Anderson et al. (1989). Plagioclase-hosted glass inclusions from white rhyodacitic pumices and the olive gray silty ash, are similar to matrix glasses from the same pumices when compared on an anhydrous basis (Fig. 6). Compositions of glass inclusions in rhyodacitic samples generally plot directly over the fields defined by the matrix glasses from the same sample (Fig. 61, and do not exhibit host crystallization enrichment effects. White rhyodacitic pumice samples also contain five plagioclase-hosted glass inclusions of andesitic

composition (61.6-62.6 SiO, anhydrous, Table 4) which are less evolved than the whole-rock and matrix glass compositions for rhyodacitic pumices and obsidians. Four mafic dacitic glass inclusions (63.8-65.8 wt.% SiO, anhydrous) have been analyzed in rhyodacitic hosts. These inclusions are more mafic than the rhyodacitic bulk rock and matrix glass compositions and also more mafic than gray dacite bulk compositions (Fig. 6). Twelve rhyodacitic to rhyolitic glass inclusions with SiO, (anhydrous) ranging from 69.2 to 72.6 wt.% are similar to the rhyolitic matrix glass from rhyodacitic host samples (Table 4). White rhyodacite sample KRA-142 contains two plagioclase-hosted glass inclusions of mafic dacite composition (SiO, anhydrous 65.6 and 65.8 wt.%). Plagioclase grains hand picked from the olive silty ash layer KRA-048 contain three glass inclu-

Table 6 Representative pyroxene compositions KRA-050 cpx wrd

KRA-050 opx wrd

KRA-053A cpx

KRA053A opx

gd

gd

rim

core

rim

core

rim

core

rim

core

TiO: SiO Al,O, Fe0 ’ MnO MgO CaO Na,O V’ Cr&h Total

51.61 0.46 1.54 8.74 0.63 14.71 20.79 0.36 0.02 0.00 98.86

50.42 0.57 2.11 9.17 0.67 14.48 20.33 0.40 0.02 0.00 98.17

53.48 0.22 0.84 18.62 1.13 24.56 1.29 0.03 0.00 0.00 100.17

53.20 0.25 1.02 19.33 1.29 23.98 1.37 0.05 0.01 0.05 100.55

53.10 0.51 1.66 9.11 0.69 15.15 20.23 0.38 0.03 0.08 100.94

51.54 0.54 1.80 9.13 0.65 14.80 20.52 0.38 0.16 0.04 99.56

53.55 0.36 0.93 18.68 1.23 24.74 1.33 0.01 0.02 0.06 100.91

53.53 0.27 0.84 18.78 1.07 24.82 1.42 0.06 0.02 0.00 100.81

%Wo a %En %Fs Mg# Mg# b

44.09 43.39 12.51 77.6 77.6

44.01 43.60 12.39 77.9 77.1

2.61 68.88 28.51 70.7 70.7

2.80 67.80 29.41 69.8 69.6

42.11 43.55 14.35 75.2 76.1

2.86 69.31 27.82 71.3 70.2

(2.6) 9 955 f 15 24

(0.6) 6

(0.8) 6

(1.0) 6

2.68 69.11 28.20 71.0 71.1 (0.6) 7

Sample: Code:

Oxide (wt.%)

10

(2.2) 10 iverage temp. (“C) Number of pairs

43.62 43.42 12.96 77.0 76.1 (1.5) 8 976 k 10 30

(0.8) 6

Fe0 * = Total Fe as FeO; 1 u = one standard deviation of Mg#; n = number of electron probe analyses. a Quadrilateral components calculated by the method of Papike et al. (19741, % Wo = percent wollastonite, %En = percent enstatite, % Fs = percent ferrosilite; Mg# = 100 X [Mg/(Mg + Fe)]. b = average Mg#.

2.56

C. W. Mandeuille et al. /Journal

of’ Volcanology and Geothermal Research 74 (1996) 243-274

sions of rhyolitic composition (SiO, anhydrous of 72.5-73.71, and three andesitic glass inclusions of 62.5-63.1 wt.% SiO, (anhydrous, Table 4). Glass inclusions within the olive obsidian samples are very similar to the evolved obsidian matrix glass when compared on an anhydrous basis (Table 4; Fig. 6). Compositions of obsidian glass inclusions plot at the center and evolved end of the field defined by matrix glasses in rhyodacite (Fig. 6). All glass inclusions in gray dacite plot within the field defined by gray dacitic matrix glasses (Fig. 6). Glass inclusions in three separate samples of gray dacite (Table 4) collectively span a wider compositional range than glass inclusions in white rhyodacite and obsidian (excluding samples containing mafic dacitic and andesitic inclusions mentioned above). Gray dacitic samples contain a few glass inclusions that are more mafic than the bulk composition of gray pumices analyzed in this study, but within the range of gray pumice whole-rock analyses of Stehn (1929) and Self (1992). Glass inclusion compositions from gray dacitic pumices that plot at more silica rich compositions than the field defined by rhyodacite matrix glasses may be an artifact of recalcula-

Table 7 Representative

10. Trace-element and REE content of whole rocks and matrix glasses Trace-element concentrations in whole-rock and corresponding matrix glass samples of 1883 erupted products have been used as indicators of magma differentiation and to assess the compatible/incompatible behavior of trace elements with respect to the phenocryst assemblage. Incompatible-element variation diagrams (e.g., La, Ce, Rb, Ba, Zr, Nb and Y) delineate more fractionated rhyodacitic samples from more mafic dacite. Trace-element concentrations in white rhyodacitic whole-rock samples exhibit slight variations when plotted on incompatible-element versus incompatible-element diagrams (Fig. 7a-d) indicating that there are compositional gradients within the white rhyodacite magma (e.g., La vs. Ce, Ba vs. Zr, Y vs. Ba, Y vs. Zr, Fig. 7). Incompatible-

iron titanium oxide compositions

Sample: Code:

Oxide TiO,

tion to anhydrous basis, a factor which probably extends the field of gray dacite matrix glasses to more evolved compositions.

KRA-045 00

KRA-052 wrd

KRA-129 wrd

KRA-053C gd

tmt. ’

ilm. h

tmt.

ilm.

tmt.

ilm.

10.50 2.45 79.20 0.78 2.38

40.80 0.35 51.20 0.86 3.41

IO.80 2.40 79.60 0.80 2.43

41 so 0.34 5 I .40 0.93 3.32

IO.30 2.27 79.90 0.72 2.46

41.00 0.30 51.00 0.85 3.58 _

0.03 95.34

0.00 96.62

0.05 96.08

0.05 97.54

0.00 95.65

0.00 96.73

tmt.

ilm.

f wt.70’0)

Al,O, Fe0 * MnO MgO “*O, CrzO, Total

x U\P x dm

L ‘ Av. T (“C) ’ f0,

No. of pairs

0.262

_

0.267

0.735 886 & 6 - 10.87 + 0.1

0.25 0.7U

891 kh -10.80+0.1 6

1

12.21 3.31 77.43 0.62 3.78 0.73 0.10 98.18 0.292

0.735 884 + 9 - 10.86 + 0.1 6

40.98 0.33 51.51 0.98 3.04 0.03 0.04 98.29 _ 0.743

906 + 10 -10.62iO.I 4

Fe0 * = Total Fe as FeO; no. of pairs = number of titanomagnetite ilmenite pairs analyzed. * tmt. = titanomagnetite. b dm. = ilmenite. ’ Cation assignment according to Stormer (1983). ’ Av. T PC) = average temperature in “C and log .f;,? calculated according to Andersen and Lindsley (1988);

C.W. Mandeville et al. / Joumal of Volcanology and Geothermal Research 74 (19%) 243-274

element concentrations in white rhyodacite wholerock and glass samples are presented in Tables 3 and 4, respectively. Two lowermost rhyodacitic pumice fall samples and two olive obsidian samples have the most evolved glass compositions on almost all incompatible-element diagrams (Fig. 7). In all samples, Y exhibits incompatible behavior and shows strong positive correlations with La, Ba, Zr, Rb, and Rb/Sr, consistent with the absence of amphibole, biotite, sphene, zircon and garnet in the phenocryst assemblage (Fig. 7a-d; Pearce, 1982; Wilson, 1989). Strontium concentrations are higher in whole-rock samples than in corresponding matrix glasses, consis-

o white rhyodacite 0 olive obsidian A gray dacite

+

1 62

64

error

“a

66

66

70

72

74

76

Weight % SIO,

+ error 64

66

66

70

72

74

72

74

Weight % SiO,

62

64

66

66

70

76

Weight % SIO, Fig. 5. Major-element variation diagrams for matrix glass samples shown in Table 4: (a) FeO’ versus SiO,; (b) CaO versus SiO,, (c) K,O versus SiO,. Error bars shown indicate 1~ analytical error on replicate microprobe analyses. Solid line denotes incremental fractional crystallization model for matrix glasses shown in Table 8, all data recalculated to volatile-free oxide sums of 100%.

257

tent with control by plagioclase (Tables 3 and 4). There is very good separation of whole-rock and matrix glass data points on the trace-element diagrams, consistent with the highly incompatible (D < 1) behavior of these elements in intermediate and felsic magmas (Arth, 1976; Nash and Crecraft, 1985; Smith and Leeman, 1987, 1993). Nb/Y ratios in the white rhyodacitic samples range from 0.1 to 0.2 and are typical of the low values found in &c-alkaline rocks (Floyd and Winchester, 1975; Gilbert et al., 1994) and within the 0.1 to 0.2 range for Krakatau samples reported by Pearce (1982). REE patterns of matrix glass from rhyodacitic samples exhibit LREE enrichment (La/Sm, 2.302.471, and relatively flat, but enriched HREE patterns, Tb/Yb, = 1.07-1.13, Yb, = 32.6-38.8 (Table 8; Fig. 8). Rhyodacitic matrix glasses show moderate negative Eu anomalies with Eu/Eu* values of 0.64 to 0.67. Rhyodacitic sample KRA-141 which contains andesitic matrix glass shows the lowest REE concentrations among rhyodacitic samples and a pattern that is parallel to those of other glasses (Fig. 8). A REE pattern of olive obsidian (Fig. 8) is parallel to those of the white rhyodacitic glasses, but has the highest REE concentrations observed for all 1883 samples and has a negative Eu anomaly (Eu/Eu* = 0.71). Olive obsidian matrix glass exhibits LREE enrichment (La/Sm, = 2.331, and a flat, but enriched (38-43 X chondritic) pattern over the HREE with Tb/Yb, = 1.12. Matrix glass from gray dacite exhibits LREE enrichment similar to rhyodacite and obsidian (La/Sm, = 2.381, but generally lower abundance values for most elements (Fig. 8). HREE concentrations in gray dacitic glass range from 35 to 42 times chondritic with a relatively flat pattern (Tb/Yb, = 1.17). Gray dacite exhibits a negative Eu anomaly (Eu/Eu * = 0.691, and only obsidian shows a relatively smaller Eu anomaly. REE contents of gray dacitic glass are lowest for all REE (Fig. 8) among 1883 samples (excluding KRA-141). Whole-rock trace-element contents of olive obsidians are similar to those in white rhyodacitic samples although at the high end of the range, and exceed the range for La and Ce in rhyodacitic whole-rock samples (Table 3). Nb/Y ratios are within the range of rhyodacitic pumices. Consequently, whole-rock data

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C. W. Mandeuille et al. / Journal of Volcano1og.v and Geothermal Research 74 (1996) 243-274

of only a couple ppm or less (Table 4). Concentrations of Cr, Ni and Co are highest in matrix glass from rhyodacitic sample KRA-141, which contains andesitic matrix glass. Cr, Ni and Co concentrations in olive obsidian matrix glass are within the range of

for obsidians plot at the most evolved end of trend lines in incompatible trace-element diagrams (Fig. 7a-d). Concentrations of compatible elements Cr, Ni and Co in rhyodacitic matrix glass span a narrow range

Field of olive obsidian

white rhyodacite O glass inclusions q

olive obsidian glass inclusions

* gray dacite glass inclusions

60

62

64

66

68

70

72

74

i

Weight % Si02 (b) 6 5

Field of gray dacite

om

Field of white

p 53

$2 1

Weight % Si02 Cc) 5

Field of white rhyodacite

Field of olive obsidian

0

1-l’. 60

1.1

62

64

5,

66

s

68

1’

70

*

72

’

8

74

73

Weight %SiO, Fig. 6. Major-element variation diagrams for glass inclusions in 1883 samples shown in Table 4: (a) Fe0 * versusSiO,; (b) CaO versus SiO,; (cl K,O versus SiO,. Fields for olive obsidian, rhyodacite and gray dacite matrix glasses determined in this study. Error bars shown indicate 1 (r analytical error on replicate microprobe analyses, all data recalculated to volatile-free oxide sums of 100%.

C. W. Mandeville

et al. / Journal of Volcanology and Geothermal Research 74 (1996) 243-274

259

I”

64 /

E B >

50 I

10=-v-. 20 25

30

..I I 35 40

.I 55

60

65

70

75

90 190

Ce mm

a

7

190 C

.L.. 45 50

210

230

b

,

,

,

,

,

210

230

250

270

290

Zrwm

250

270

290

Zr ppm

340 d

350

360

370

390

390

400

410

Bawm

Fig. 7. Incompatible-element variation diagrams for whole-rock samples (open symbols) and matrix glasses (solid symbols): (a) La versus Ce plot of whole-rock and matrix glass samples. * denotes pre-1883 andesite sample RR.4076, see Table 3; (b) Y versus Zr; (c) Rb/Sr versus Zr; (d) Rb/Sr versus Ba, * same as in (a). Error bars shown indicate 1(T analytical error on replicate XRF analyses of reference standards. Shown for reference in (a) are whole-rock data from Anak Krakatau samples spanning from 1928 to 1963 of Nishimura et al. (19801, a whole-rock analysis of 1975 Anak Krakatau basaltic andesite lava (this study) and whole-rock analysis of a pre-1883 andesitic dike from Rakata (this study, Table 3), and whole-rock analysis of Sukadana plateau basalt from southern Sumatra (Nishimura et al., 1986).

rhyodacitic samples (Table 4). Matrix glass from gray dacite has higher Cr and Co concentrations than white rhyodacitic samples (excluding KRA- 14 11, but slightly lower Ni content than the rhyodacite and obsidian matrix glasses (Table 4). However, Ni concentration of one gray dacite sample (19 ppm) is equal to that observed in pre-1883 andesite (Table 3).

11. Dissolved volatile content of the 1883 magma and degassing to the atmosphere The petrologic method of estimating volatile discharge from explosive eruptions is the only available approach for eruptions prior to advent of remote sensing methods, and/or where ice core records are

not available (Anderson, 1974; Devine et al., 1984; Anderson et al., 1989; Sigurdsson, 199Oa). This method scales the volatile concentration difference observed between silicate glass inclusions (trapped in phenocrysts) and matrix glass to the mass of the erupted glass in order to estimate the total volatile degassing. The presence of an excess vapor phase in the chamber, magma degassing prior to eruption, and degassing from non-erupted magma remaining at depth in the chamber are not accounted for by the method. Consequently, current petrologic estimates of degassing are minimum estimates (Devine et al., 1984; Sigurdsson et al., 1990a; Westrich and Gerlath, 1992; Gerlach et al., 1994). Volatiles by difference from 51 analyses of white rhyodacite-hosted glass inclusions and 4 analyses of obsidian-hosted glass inclusions ranges from 3.0 to

C. W. Mandeli&

260

Table 8 Rare earth element concentrations

et al. / Journal of Volcanology and Geothermal Research 74 (1996) 243-274

of glasses and whole rocks

Sample: Code b:

045-2 mg 00

0.50 mg wrd

OS2 mg wrd

053A mg gd

129 mg wrd

141 mg wrd

076 wr a a

025 wr ba

@pm) ’ La Ce Pr Nd Sm Eli Gd Tb DY Ho Er Tm Yb Lu

32.3 70.0 9.2 38.8 9.0 2.10 8.9 1.61 9.8 2.2 1 6.4 1.01 6.6 0.99

32.6 69.3 9.4 38.1 9.1 I .97 9.5 1.60 9.5 2.15 6.6 1.01 6.4 0.99

31.7 68.7 9.0 36. I 8.7 1.96 9.2 I .56

31.6 67.6 x.9 36.6 8.6 1.96 8.7 I .57 9.0 1.95 5.8 0.9 1 6. I 0.91

3 I .4 68.4 8.9 36.6 8.4 1.87 8.6 1.49 9.1 1.97 6.1 0.94 6.3 0.94

30.6 64.7 8.4 34.1 8.0 1.76 7.9 1.31 8.1 1.73 5.2 0.85 5.6 0.80

24.7 52.7 6.3 27.8 7.2 2.04 6.8 1.20 7.0 1.39 4.3 0.65 4.5 0.74

15.1 33.6 4.2 19.3 5.4 1.74 5.3 0.97 5.2 1.02 3.1 0.46 3.3 0.49

9.4 2.09 6.4 0.97 6.6 0.98

d wr = whole rock. ’ Codes same as in Table 3 and Table 4, ba = basaltic andesite. ’ ICP-MS analyses on glasses at Harvard University (R. Murray, analyst), ICP-MS analyses on whole rocks at Geological (C. Mandeville, analyst).

-+-

~~ KRA-045 glass; oo KRA-050 glass; wrd ~~~KRA-052 glass; wrd

~~~~ KRA-129 glass; ~~ KRA-053 glass; *--KRA-141 glass; ~ KRA-076 whole - - - - - KRA-025 whole Anak 179 whole

La

Ce

Pr

Nd

Pm Sm

Eu

Gd

Tb

Survey of Japan

Dy

Ho

Er

wrd gd wrd rock; a rock; ba rock; ba

Tm

Yb

Lu

Element Fig. 8. Chondrite-normalized REE plot of Krakatau glass separates, and whole-rock data from pre-1883 andesite KRA-076 and Anak Krakatau basaltic andesite KRA-025. Note parallelism of patterns from whole-rock samples with 54 to 64 wt.% SiO, to 1883 pure glass separates (wt.% SiO, > 691. Codes same as in Table 3 and Table 4, ba = basaltic andesite. Chondrite normalization factors from Sun and McDonough (1989). Sample 179 is a whole-rock IWE analysis of basal Anak Krakatau basaltic andesite ash from Nishimura et al. (1980).

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C. W. Mandeville et al. / Journal of Volcanology and Geothermal Research 74 (1996) 243-274

1883 erupted products. Stabilization of amphibole in rhyodacitic bulk compositions determined from experimental studies and in natural assemblages requires from 4 to 6 wt.% H,O (Merzbacher and Eggler, 1984; Rutherford and Devine, 1988; Drum and Bacon, 1989; Geschwind and Rutherford, 1992; Rutherford, 1993; Gardner et al., 1995). Average volatiles by difference in mafic dacitic inclusions within rhyodacitic hosts is 4.5 wt.% (la = 0.82, n = 3). Sulfur concentration in rhyodacitic glass inclusions ranges from 108 to 395 ppm with an average of 200 ppm (Table 9). Chlorine concentrations in rhyodacitic glass inclusions range from 2064 to 3679 ppm with an average of 2722 ppm (Table 9). Volatiles by difference in glass inclusions within gray dacite ranges from 2.1 to 6.4 wt.%. Average volatiles by difference for gray dacitic samples is 4.8 wt.% (Tables 4 and 91, and implies a pre-eruption

6.8 wt.% with an average of 4.5 wt.% (lo = 0.92, n = 55). It is assumed that most of the volatile component in these rhyolitic glasses consists of H,O due to the much greater solubility of water in rhyolitic glasses relative to CO, (Silver et al., 1990; Fogel and Rutherford, 1990). A recent comparison of water determination techniques in hydrous rhyolitic glasses (Devine et al., 1995) indicated that volatiles by difference from electron microprobe analyses correlate well with H,O contents measured by FTIR and ion microprobe, though the accuracy and precision of the electron microprobe analyses is typically greater than 0.5 wt.%. This indicates that the actual water content of glass inclusions in Krakatau 1883 erupted products is likely to be in the 3-4.5 wt.% range, and implies an average water content in rhyodacitic glass inclusions of 4.0 f 0.5 wt.%. This estimate is consistent with the absence of amphibole in

Table 9 Petrologic estimates of water, sulfur and chlorine degassed from Krakatau 1883 magmas Magma type

H,O COW.Diff. a (wt.%)

Mass H,O (g)

S cont. diff. b (eem)

Mass sulfur (g)

Cl cont. diff. ’ (eem)

Mass chlorine @

Rhyodacite Gray dacite Andes&e Total (this study)

3.4 3.6 3.9

8.7 x 1o14d 3.9 x 10’3 e 1.1 x 1Or3f 9.2 X 1014

103 103 324

2.6 x 1.1 x 1.0 x 2.8 x

lOI 10” 10” lOI

351 428 810

9.0 x 5.0 x 2.3 x 9.7 x

9.6 x 10”

152

3.6 x lOI

Devine et al. (1984) Ice core records Hammer et al. (1980) Legrand and Delmas (1987) Zielinski (1995)

40

10’2 10” 10” 10’2

1.8 x lOI 1.0-1.2 x 10’3 1.9 x 10’3

Densities calculated according to Bottinga et al. (1982). a H,O cont. diff. = average dissolved volatile concentration (by difference) in rhyodacitic glass inclusions (4.5 wt.%, lu= 0.92, n = 55) - average dissolved volatile concentration (by difference) in rhyodacitic matrix glass (1.1 wt.%, 1(T= 0.96, n = 46, assumed to be H,O), average dissolved volatile concentration (by difference) in gray &cite glass inclusions (4.8 wt.%,, 1tr = 1.0, n = 26) - average dissolved volatile concentration (by difference) in gray dacite matrix glass (1.2 wt.%, 1u = 0.5, n = 8, assumed to be H,O). b Average dissolved sulfur concentration in rhyodacite glass inclusions (200 ppm, 1u = 70, n = 41) - average dissolved sulfur concentration in matrix glass (97 ppm, 1(T= 41, n = 48) average dissolved sulfur concentration in gray dacite glass inclusions (234 ppm, 1cr = 78, n = 16) - average dissolved sulfur concentration in matrix glass (131 ppm, 1~ = 46, n = 11) average dissolved sulfur concentration in andesitic glass inclusions (879 ppm, 1u = 267, n = 3) - average dissolved sulfur concentration in matrix glass (555 ppm, n = 2). ’ Average dissolved chlorine concentration in rhyodacitic glass inclusions (2722 ppm, 1u = 338, n = 57) - average dissolved chlorine concentration in matrix glass (2371 ppm, 1u = 369 , n = 57 ), average dissolved chlorine concentration in gray dacite glass inclusions (2748 ppm, 1(T= 521, n = 16) - average dissolved chlorine concentration in matrix glass (2320 ppm, 1cr = 400, n = 12). average dissolved chlorine concentration in andesitic glass inclusions (2689 ppm, 1(T= 1468, n = 4) - average dissolved chlorine concentration in matrix lass (1879 ppm, n = 2). f 90% of 12.5 km3 (DRE) corrected for 10% crystallinity, rhyodacite density = 2400 kg/m3. ’ 4% of 12.5 km3 (DRE) corrected for 10% crystalhnity, gray dacite density = 2410 kg/m3. f 1% of 12.5 km3 (DE) corrected for 10% crystallinity, andesite density = 2550 kg/m3.

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of Volcanology and Geothermal Research 74 (19961243-274

dissolved water concentration in gray dacitic magma of 4.8 t_ 1.0 wt.% similar to the concentration in rhyodacite. Sulfur concentrations in dacitic glass inclusions are higher than in rhyodacitic inclusions and range from 147 to 376 ppm, with an average pre-eruption sulfur concentration of 234 ppm (Table 9). Chlorine content of gray dacite-hosted glass inclusions ranges from 2428 ppm to 3754 ppm with an average of 2748 ppm (Table 9). Rare andesitic glass inclusions found in two rhyodacitic samples contain from 700 to 1186 ppm sulfur with an average concentration of 879 ppm (Table 9). Dissolved chlorine concentrations in these andesitic inclusions range from 1824 to 5225 ppm with an average concentration of 2690 ppm (Table 9). Estimated water discharge from the erupted volume of rhyodacite (including obsidian) is 8.7 X 10“’ g based on an average concentration difference of 3.4 wt.% volatiles by difference (assumed to be mostly water) between glass inclusions (4.5 wt.%) and matrix glasses with average VBD of 1.1 wt.% (n = 46, I CT= 0.96 Table 4, Table 9). Dissolved volatile concentration in degassed rhyodacitic matrix glasses (by difference method) range from 0.0 to 2.4 wt.%. Volatiles by difference in obsidian matrix glasses are generally less than 1.0 wt.% with an average of 0.9 wt.% (1 g = 0.5 1. n = 24. Table 4). Sulfur degassing from erupted rhyodacite (including rhyodacitic obsidian) is 2.6 X 10” g S (Table 9). Chlorine discharge from rhyodacite is 9.0 X IO” g based on average melt inclusion concentration 2722 ppm and matrix glass with average Cl concentration of 2371 ppm (Table 9). Dissolved volatiles in gray dacitic matrix glasses (by difference) are generally higher than in rhyodacite with an average of 1.2 wt.% (Tables 4 and 9). Water discharged from the erupted volume of gray dacite is 3.9 X lOI g (Table 9). Estimated sulfur degassed from the gray dacite magma is 1.1 X 10’ ’ g, and estimated chlorine discharge is 5.0 X 10” g (Table 9). Though we have documented the occurrence of andesitic liquid compositions in 1883 erupted products, the relative amount of andesitic material erupted is still highly uncertain. We have found very minor amounts of andesitic glass in only two out of eleven samples and generously assign a value of 1% to the

estimated amount of juvenile andesite in 1883 erupted products. Estimated water discharge from erupted andesite is 1.1 X 1013 g H,O based on a concentration difference of 3.9 wt.% but this is highly uncertain due to only two andesitic matrix glass analyses (Table 9). Estimated sulfur discharge from erupted andesite is 1.0 X 10” g, and chlorine discharge from erupted andesite is estimated at 2.3 X 10” g (Table 9). Total magmatic water discharge from degassing of erupted magma is estimated at 9.2 X lOI g H,O (94.5% from rhyodacite, 4.5% from gray dacite, and 1% from andesite). Total sulfur discharge is 2.8 X 10” g S (93% from rhyodacite, 4% from gray dacite, and 3% from andesite). Total chlorine discharge is estimated at 9.7 X lOI g Cl (92% from rhyodacite, 5% from gray dacite, and 3% from andesite). Our petrologic estimate of sulfur discharge from degassed 1883 magma is approximately three times larger than that of Devine et al. (1984) (9.6 X 10” g S). This is due to three factors: (1) we are using a new volume estimate for the eruption of 12.5 km”: (2) we have accounted for the shift in sulfur Ka X-ray wavelength due to oxidation state, 3) we have accounted for contributions from more mafic gray dacite and andesite. Of these factors, the shift in sulfur K CYX-ray wavelength may be most important as the average sulfur concentration difference observed between glass inclusions and matrix glasses in rhyodacitic magma (103 ppm) is roughly 2.6 times the 40 ppm concentration difference observed by Devine et al. (1984). The eleven analyses of Devine et al. (1984) could also have been on glass inclusions at the low end of the concentration range defined by this study. Approximately 78% of the tephra generated during the 1883 eruption was deposited in the marine environment (Mandeville et al., 1996) mostly in the form of pyroclastic flows. This process could have vaporized considerable amounts of seawater due to the relatively high ( > 500°C) emplacement temperature of the pyroclastic flows as much as 12 km from source (Mandeville et al., 1994). The heat output from one half of the erupted mass cooling from 500 to 20°C is 1.1 X lOI J (1.4 X 1013 kg; latent heat 3 X lo5 J/kg; specific heat lo3 J kg-’ ‘C’). Assuming fully efficient energy transfer and 2.6 X 10h

C. W. Mandeville et al. / Journal of Volcanology and Geothennal Research 74 (1996) 243-274

J/kg energy consumption in seawater vaporization (Sigurdsson et al., 199Ob), this energy source could vaporize 4.2 X 10” kg of seawater. Using a mean sulfate concentration in seawater of 2.7 g/kg (salinity = 35 parts per thousand), and a mean chlorine concentration of 19.4 g/kg, evaporation of this mass of seawater could have contributed 3.8 X lo’* g S, roughly equivalent to the amount estimated from degassed magma above, and 8.1 X lOi g Cl, or an order of magnitude more Cl than the estimated amount from degassed Krakatau magma. Much of this inferred S and Cl discharge could have been injected to stratospheric height by secondary coignimbrite ash columns as discharge rate most likely exceeded lo8 kg/s during the flow phase of the eruption (Woods and Wohletz, 1991). 12. Pressure and temperature constraints on 1883 magma Constraints on pre-eruptive temperature, total pressure, and fo, conditions in the 1883 magma

263

chamber have been determined from mineral-melt equilibria, mineral Fe-Mg exchange equilibria, volatile content of plagioclase-hosted glass inclusions, and the phenocryst assemblage of 1883 erupted products. Iron-titanium oxide temperatures calculated for five rhyodacitic samples and one obsidian range from 876 to 897°C with an average temperature for all titanomagnetite-ilmenite pairs of 887°C (1 rr = 7.O”C, IZ= 26) using the calibration of Andersen and Lindsley (1988) and the structural model of Stormer (1983). All titanomagnetite-ilmenite pairs were touching and in contact with matrix glass and passed the Mg/Mn equilibrium test of Bacon and Hirschmann (1988). Furthermore, temperature variation within the rhyodacitic magma was probably less than 10°C as five out of six rhyodacitic samples have Fe-Ti oxide temperatures between 880 and 890°C. Oxygen fugacity determined from Fe-Ti oxides in white rhyodacitic samples ranges from log fo, of - 10.79 to - 11.07 (Fig. 9). These log fo,values agree well with the oxidation state of the magma as

-9

(Dtuitt and Bacon, 1999) 0

-11

s

+

1883 graydacite

-13 800

850

900

950

1000

Temperature OC Fig. 9. Temperature versus log fo, plot illustrating field of Krakatau 1883 samples and fields for Mt. Mamma Climactic pumice and Krakatau Pre-1883 andesite; MN0 and FMQ buffer curves from Chou (1978), NNO buffer curve from Huebner and Sato (1970). Dashed line shows transition zone of sulfur dissolved as sulfate in the melt vs. sulfide based on experimental work on stability of anhydrite in El Chichon trachyandesite (Luhr, 1990).

264

C. W. Mandeuille et al./Journal

of Volcanology and Geothermal Research 74 (1996) 243-274

determined from the shift (AASKa = 2.45~10~ A, relative to FeS) observed in SK (Y X-ray wavelengths of glass inclusions (Fig. lo), which suggests approximately 2 log units above the QFM buffer (Chou, 1978) and 80% of sulfur in the glass inclusions as sulfate (Carroll and Rutherford, 1988). Iron titanium oxide temperatures in gray dacitic pumices are slightly higher than in rhyodacite and range from 887 to 913°C. Log fo, ranges from - 10.6 to - 10.9, although the titanbmagnetite and ilmenite in one sample (KRA-053C) were not in contact (Table 7). The slightly higher (IO-20°C) temperature of the gray dacitic samples relative to rhyodacite has been confirmed using the clinopyroxene-orthopyroxene Fe-Mg exchange thermometer of Wood and Banno (1973). Twenty four clinopyroxne-orthopyroxene rim composition pairs in white rhyodacite yield an average temperature of 955 + 15”C, and thirty rim pairs in gray dacite yield an average temperature of 976 + 10°C (Table 6). All pyroxene compositions used for thermometry had over 90.8% or more quadrilateral components, and passed the stoichiometric criteria of Papike et al. (1974). Temperatures from Fe-Ti oxides are more likely to represent the last temperature of equilibration prior to quenching, because subsolidus equilibria in the pyroxenes commonly lock in at higher temperatures than the oxide phases, and temperatures deter-

700

s 8 7 3j E 2 :

600

500

0 400

r

I

0.6130

0.6135

-

KRK-053C

Sko counts

-

KRK-053A

Ska counts

0.6140

0.6145

0.6150

Sin 8

Fig. 10. Sulfur Ka X-ray peak search on Krakatau glass inclusions within host samples KRA-053C, and KRA-053A according to Carroll and Rutherford (1988). Shown for reference are sin 0 peak positions for sulfur dissolved in melt as sulfide (S*- ) and sulfate (S6 + ).

4 kb

0

10 20 30 40 50 60 70 80 90 100

Anorthite (mol%) Fig. I 1. Plagioclase liquidus/solidus diagram illustrating the nature of plagioclase liquidus/solidus surfaces in the binary An-Ab system (Bowen, 1913), Ab-An-H,0 system (Yoder et al., 19571, the Ab-An-Di system (Morse, 19801, and in natural multi-component systems under total pressure of 1, 2 and 4 kbar from experimental data of Housh and Luhr (1991) (references therein). Multi-component glass compositions are plotted in terms of normative albite and anorthite components [ 100 X An/(An + Ab)]. 0 denote 1 and 2 kbar experimental data of Housh and Luhr (1991). 0 denote Krakatau 1883 rhyodacite sample pairs (mode in plagioclase rims, and average matrix glass). A denote Krakatau 1883 gray dacitic sample pair. 0 denote experimental data of Rutherford et al. (1985) at 1.8-2.3 kbar.

mined from Fe-Ti oxide phases are closer to experimental temperatures on dacitic to rhyolitic bulk compositions (Rutherford and Devine, 1988; Geschwind and Rutherford, 1992; Frost and Lindsley, 1992). Pre-eruption total pressure for white rhyodacitic magma is derived from comparison of natural plagioclase rim and matrix glass compositions to experimental determinations of plagioclase-liquid equilibria of Housh and Luhr (1991). Average matrix glass compositions in white rhyodacite and modes in plagioclase rim compositions plot near or just to the right of the 1 kbar experimental data of Housh and Luhr (1991) (see Fig. 11). Matrix glass compositions

C. W. Mandeville et al. / Journal of Volcanology and Geothemuzl Research 74 (1996) 243-274

and modes in plagioclase rim compositions from gray dacite plot to the right of the data from white rhyodacite, suggesting slightly higher total pressures (Fig. 11). Additional constraints on total pressure and PHzO can be derived from the absence of amphibole in the phenocryst assemblage, and the absence of any amphibole breakdown products (clots of plagioclase, clinopyroxene and titanomagnetite, see Rutherford and Hill, 1993) within Krakatau 1883 erupted products, and HREE concentrations observed in rhyodacite and gray dacite matrix glasses. Volatiles by difference in rhyodacitic glass inclusions average 4.0 + 0.5 wt.% and are near the saturation concentration for rhyolitic melts at 1.O to 1.5 kbar total pressure (Silver et al., 1990). Absence of amphibole phenocrysts probably reflects the fact that the rhyodacitic magma achieved water saturation at pressures less than the lower stability limit of amphibole under water-saturated conditions (1.5-1.6 kbar at 89O”C, NNO + 1 and 1.4 kbar at 850°C; Geschwind and Rutherford, 1992; Gardner et al., 1995). This implies water saturation at relatively shallow levels in the crust of 4-5 km (assuming a density of 2.55 g/cm3). Average chlorine content of glass inclusions in Krakatau rhyodacite is 2720 f 338 ppm and very close to the experimentally determined saturation value of 2790 for natural rhyolitic melts at 1.0 kbar total pressure and 830-850°C (Metrich and Rutherford, 1992). This provides independent confirmation of shallow level staging of magma and vapor saturation just prior to eruption.

265

andesitic and rhyodacitic compositions (Anderson, 1976; Bacon, 1986). Andesitic glass inclusions and matrix glasses within two rhyodacitic hosts indicates the presence of andesitic magma in the 1883 magma chamber (Table 4; Figs. 5 and 6). Presence of rhyolitic glass inclusions within macroscopically homogenous gray dacitic pumices and isolated patches of white matrix glass indicates that phenocrysts within gray dacite were in contact with rhyodacitic to rhyolitic liquids and that mixing between liquid compositions occurred in the chamber. Anorthite-rich plagioclase rim compositions ( > An 55_6,,)both in rhyodacite and gray dacite, indicate disequilibrium between a fraction of the phenocryst populations and host liquids in these low crystallinity rocks. Similar modes for plagioclase rim compositions from rhyodacite and gray dacite (Fig. 4) and considerable overlap of plagioclase core compositions is evidence for extensive mixing in the chamber as opposed to simple syneruptive mingling (Druitt and Bacon, 1989). Reversely zoned phenocrysts of orthopyroxene in gray dacite and rhyodacite, and slightly higher pyroxene temperatures in gray dacite indicate contact with/or derivation from higher temperature, andesitic magma. Presence of andesitic glass inclusions and matrix glass within rhyodacite requires that the two compositions were once in contact, and that andesitic liquids and associated phenocrysts were entrained into a much greater volume of rhyodacite.

14. Origin of compositional diversity within 1883 magma 13. Evidence for magma mixing Evidence for magma mixing in 1883 erupted products consists of compositionally banded pumices, heterogeneous matrix glass composition on a thinsection scale, disequilibrium phenocrysts in gray dacitic and rhyodacitic pumices, glass inclusion and matrix glass compositions that are less evolved than the whole-rock and average matrix glass compositions of their hosts (Anderson, 1976; Bacon, 1986; Sigurdsson et al., 199Oa). These features and the more variable composition of matrix glass in gray dacitic pumices indicates magma mixing and the generation of hybrid pumices intermediate between

The entire suite of glass inclusions and matrix glasses from the 1883 products produce well defined linear trends on major-element variation diagrams (Figs. 5 and 6). Many of these compositions are clearly hybrid, resulting from mixing or comingling of mafic and evolved magma. Liquid compositions present in 1883 erupted products, as represented by matrix glasses and glass inclusions, range from andesite (61.6% SiO,) to dacite (70.0% SiO,) and rhyolite (74.4% SiO,). In order to evaluate the relationships between the observed end members in the 1883 compositional spectrum the major-element data have been modeled

C. W. Mandeville et al. /Journal

266 Table 10 Incremental fractional matrix glasses

crystallization

model

of Volcanology and Geothermal Research 74 (19961243-274

of 1883 Krakatau

Step 1

Step2

Step3

Step4

Step5

Step6

63.18 1.13 15.52 6.00 0.16 2.26 4.78 5.53 1.43

65.76 0.94 15.51 4.61 0.20

68.25 0.80 14.74 3.60 0.19 1.27 3.38 5.75 2.02 0.016

70.82 0.68 14.30 2.78 0.12 0.81 2.46 5.56 2.47 0.072

72.71 0.52 14.12 2.35 0.04 0.60 I .79 5.47 2.40 0.145

74.44 0.44 13.41 I .72 0.11 0.36 1.34 5.64 2.54 0.015

7.1 0.8 2.4 0.9 0.2 88.6

3.8 0.0 1.8 0.3 0.4 93.7

5.1 0.9 0.0 0.7 0.1 92.5

32.2

36.5

41.3

SiO, TiO Al,& Fe0 * MnO MgO CaO

Na,O K,O R2

1.57 4.09 5.46 1.86 0.299

Cumulative crystallizing phases Plag 6.6 8.9 OPx 1.7 1.0 CPX 2.3 1.1 Tmt 1.7 1.4 Ilm 0.2 0.2 Liq. % 100.0 87.5 87.4 % of initial liquid mass crystallized

12.5

23.5

Fe0 * = Total Fe as FeO; R2 = sum of residuals squared; Plag = plagioclase; Opx = orthopyroxene; Cpx = clinopyroxene; Tmt = titanomagnetite; Im = ilmenite; Liq. % = percent of liquid remaining after each crystallization step. % of initial liquid mass crystallized determined from percent of liquid remaining after each crystallization step.

by incremental fractional crystallization of an andesite to produce rhyodacite (Table 10; Fig. 5). The incremental fractional crystallization model makes use of five matrix glass analyses from 1883 samples beginning with andesite (Step 1; Table 10; Fig. 5) to the most evolved rhyolite (Step 6; Table 10; Fig. 5). In all crystallization steps sums of squared residuals were very low to near zero, indicating very good to excellent solutions (Table 10; Fig. 5). In each step plagioclase was the most abundant phenocryst phase extracted in the crystallizing assemblage, consistent with modal data in 1883 products, and in pre-1883 andesite. Amounts of plagioclase extracted decrease slightly from steps 3 to 5, and increase again in the last crystallization step. Mass balance solutions require removal of titanomagnetite and ilmenite in all crystallization steps. Extraction of clinopyroxene is required in all but the last crystallization step, and

orthopyroxene must be extracted in all steps except number 5 (Table 10). This model implies approximately 41% crystallization of an andesitic parent magma to produce the most evolved rhyolitic matrix glass. An alternative incremental fractional crystallization model, which excludes the step 3 point (a matrix glass in gray dacite) and directly crosses an apparent compositional gap from step 2 to step 4 involves 22.5% crystallization (15.1% plagioclase, 1.8% orthopyroxene, 3.2% clinopyroxene, 2.2% titanomagnetite and 0.3% ilmenite), has sum of squared residuals CR’> of 0.021 and also implies 41% total crystallization of the andesite to produce rhyolite. If the majority of gray dacite compositions are considered to be hybrid compositions, then there is suggestion of a compositional gap from 65.8 to 69.1% SiO, in liquid compositions defined by the glass inclusion data (Fig. 6) and matrix glasses (Fig. 5, excluding data from hybrid gray dacite pumices) This gap may be due to the shallow slope of the liquidus surface in evolved calcalkaline compositions crystallizing plagioclase, orthopyroxene, clinopyroxene, Fe-Ti oxides, + hornblende thus allowing for large changes in liquid composition (> 11 wt.% SiO,) over relatively small temperature drops of approximately 20°C and 20-40% crystallization (Hildreth, 1981; Rutherford et al., 1985; Grove and Donelly-Nolan, 1986; Bacon and Druitt, 1988; Sigurdsson et al., 1990a). The major-element fractional crystallization model is supported by trace-element and REE data from a pre- 1883 andesitic dike, a 1975 basaltic andesite lava flow from Anak Krakatau (Table 11, and data from Nishimura et al. (1980) on early Anak Krakatau basaltic andesite ash (Figs. 7a and 8). A pre-1883 andesitic lava flow from Panjang has a “Sr/ 86Sr ratio of 0.70434 + 0.00018 (Whitford, 1975) that is within analytical error of the “Sr/ 86Sr ratio of 1883 dacite (0.7044, Sigmarsson, 1990). Sigmarsson (1990) proposed that the Krakatau 1883 dacite is a result of approximately 75% crystallization of primitive Krakatau basalts, on the basis of U-Th isotopic systematics. The REE patterns for a basaltic andesite lava from Anak Krakatau and a pre-1883 andesite are approximately parallel to the patterns of the dacitic and rhyodacitic glasses (Table 8; Fig. 8) despite having

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lower REE concentrations, indicating that Krakatau basaltic andesites to andesites are most likely parental to the 1883 rhyodacitic magma. Ratios of light to heavy REE remain relatively constant from basaltic andesite whole rock (La/Yb, of 3.3-3.5) through dacitic and rhyodacitic glass samples (La/Yb, 3.53.9) consistent with derivation of 1883 dacite and rhyodacite through crystal fractionation of basaltic andesites-andesite. The REE pattern of the basaltic andesite is LREE enriched, with La/Sm, of 1.75, though it is less enriched than the dacites-rhyolites, and exhibits a relatively flat HREE pattern with Tb/Yb, of 1.29. Slopes of LREE patterns in medium-K talc-alkaline rocks commonly increase in steepness with increasing silica (Gill, 1981) and the slightly lower La/Sm, value of the basaltic andesite sample from Anak is consistent with its lower degree of differentiation. La/e ratios of Anak Krakatau basaltic andesite ash and lavas and the pre-1883 andesite are similar to ratios observed in 1883 rhyodacite and gray dacite samples (Fig. 7a), consistent with fractionation of andesitic to rhyodacitic magmas from parental Krakatau basaltic to basaltic andesite magma deeper in the Krakatau magma system. It is apparent from the relatively flat but enriched (20-40 X chondritic) and subparallel HREE patterns of 1883 dacite-rhyodacite, pre-1883 andesite, and Anak basaltic andesite samples, that Krakatau magmas have not undergone amphibole or garnet fractionation, nor is it likely that these magmas represent melts derived from amphibole or garnet-bearing lower crust (Gill, 1981; Wilson, 1989).

15. Discussion 15.1. Magma chamber model

A model of the magmatic system beneath Krakatau prior to the 1883 eruption must account for the following features observed or deduced from 1883 erupted products: (1) a relatively shallow chamber (4-5 km) that contained at least 12.5 km3 (DRE) of magma; (2) the existence of a large volume of relatively homogeneous rhyodacite at a temperature of 880-890°C; (3) presence of dense, and more mafic dacitic to andesitic magma at higher temperature (890-913 and 982-lOOO”C, respectively); (4)

261

occurrence of fractional crystallization to generate the evolved rhyodacitic compositions; (5) the opportunity for mixing and hybridization of mafic and evolved liquids. The magma chamber model in Fig. 12 shows a compositionally and thermally zoned magma system similar to that proposed by Sigurdsson and Sparks (1981) for the Askja chamber and the Mount Mazama chamber (Bacon and Druitt, 1988; Drum and Bacon, 1989). Rhyodacite is by far the predominant magma produced during the 1883 eruption. Fe-Ti oxide temperatures determined from rhyodacitic samples indicate a narrow temperature range of 880-890°C which suggests that the upper portion of the chamber was probably convecting and thereby relatively homogeneous in temperature and composition. Density of the rhyodacitic magma calculated from oxide data on glass inclusions with 4.0% H,O (Bottinga et al., 1982), modal composition, and mineral densities is 2.40 g/cm3. In contrast, the density of andesitic liquid calculated from glass inclusion compositions with 4.0% H,O is 2.55 g/cm3 and thus the magma chamber would have been stratified from a basal andesitic zone to an upper rhyodacitic part. Density stratification in the chamber may have originated from recharge of andesitic magma into the basal portion of a magma chamber composed largely of rhyodacite (Sigurdsson and Sparks, 1981; Bacon and Druitt, 1988; Drum and Bacon, 1989). Crystallization of plagioclase, orthopyroxene, clinopyroxene and Fe-Ti oxides, either along the floor and/or along the walls (Sparks et al., 1984; Bacon and Druitt, 1988; Druitt and Bacon, 1989) could have produced more evolved melt which could then convect upwards due to compositional buoyancy, thereby contributing more evolved melt over time and causing the rhyodacite volume to grow. The uppermost portion of the chamber may have become lined with quenched rhyodacitic obsidian, which, based on contrasting glass inclusion and matrix glass volatiles by difference (Table 41, appears to have degassed through conduit walls. Evidence for chemical zonation in the upper rhyodacitic part consists of incompatible-element enrichment of olive obsidian samples, and lower rhyodacitic pumice fall samples. A significant feature of the incompatible-element plots of Fig. 7 is that matrix glass from two of the lowermost rhyodacitic pumice fall deposits (samples KRA-050 and

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of’ Volcanology and Geothermal Research 74 (19961 243-274

had a thermal gradient. Fe-Ti oxide temperatures derived from more mafic gray dacite samples range from 890 to 913°C and are intermediate between that of rhyodacite and andesite, consistent with presence of a thermal gradient at the base of the overlying rhyodacitic magma. Density calculations for the gray dacite based on glass inclusion analyses, modal composition, plus 4.0% H,O yield a value of 2.41 g/cm3 or greater because of titanomagnetite microlites. This magma probably formed from addition

KRA-052) and olive obsidian samples plot towards the most evolved end of the incompatible trace-element trends along with matrix glass from lithic and crystal poor pyroclastic flow sample KRA-129. Matrix glasses from these obsidians and rhyodacitic pumice fall samples are the most evolved 1883 samples, and their density calculated from oxide data is 2.39 g/cm3 The interface between low density rhyodacitic and denser andesitic portions of the chamber must have

1-J Pre-1883 Rakata Basaltic Cone W

6erboewatan

&

’ /// Danan vents

Set-Jung

E sea level

Panjang

ex-:$20nduit: __ ,

Upper Pliocene

Granitic Crust Cretaceous?

Base of crust I

I

0

1

1

1

1

I

5 km

Fig. 12. Krakatau magma chamber model prior to 1883 eruption. Magma densities calculated according to Bottinga et al. (1982) (see text). Magma temperatures calculated from Fe-Ti oxides (see text). Magma volume from Mandeville et al. (1996).

C. W. Mundeville et al. / Journal of Volcanology and Geothermal Research 74 (1996) 243-274

and mixing of residual andesitic liquids that attained densities nearer to rhyodacite, thus mixing with lower portions of the overlying rhyodacite. Decreased liquid density resulting from andesite crystallization allowed for localized mixing at the base of the rhyodacite thus forming a small intermediate convection cell of hybrid magma between the rhyodacite above, and andesitic magma below (Fig. 12). Both processes may have contributed towards development of a compositional gap in 1883 liquid compositions. The recharge and convective fractionation model (Bacon and Druitt, 1988; Drum and Bacon, 1989) is attractive in that many of the pyroxene phenocrysts in 1883 gray dacite and rhyodacite magmas are reversely zoned. Subequal proportions of reversely and normally zoned plagioclase in gray dacitic samples and reversely zoned plagioclase in rhyodacite also indicate a possible recharge event. Mixing of minor amounts of intermediate andesite and large amounts of evolved rhyodacite is more likely to occur than mixing between basaltic andesite and rhyodacite as was shown by the thermal and physical property modelling of Sparks and Marshall (1986). Plagioclase-liquid equilibria and strongly unimodal plagioclase rim compositions suggest that the 1883 rhyodacitic magma staged at a total pressure of 1.0 to 1.5 kbar or roughly 4-5 km depth. A recent microseismicity survey by Harjono et al. (1989) has revealed the presence of a small seismic wave attenuation zone under Krakatau at a crustal depth of 8-9 km, and a much larger seismic attenuation zone at the base of the crust at 22 km depth beneath the volcano. The present day attenuation zone at 8-9 km depth most likely either extended to much shallower depth of 4-5 km prior to the catastrophic 1883 eruption, or was at least in communication with a shallower magma chamber, as evidenced by the 270m-deep caldera formed during the 1883 eruption (Sigurdsson et al., 1991a). Cross section A of Harjono et al. (1989, their fig. 8a) indicates a small attenuation zone at 5-6 km depth. Eruption of Anak Krakatau basaltic ash and lavas 44 years after the 1883 eruption (Stehn, 1929; Hardjadinata, 19851, was probably facilitated by extensive evacuation of crystal-poor 1883 rhyodacite from a shallow, 4- to 5-km-deep reservoir that may have acted as a density filter to Anak basaltic andesite magma.

269

Krakatau 1883 erupted products are remarkably similar to those of the 6845 + 50 yr B.P. eruption of Mount Mazama in terms of erupted compositions, temperature and fo, of the magma, estimated preeruptive dissolved H,O concentration, and total pressure (Bacon and Druitt, 1988; Druitt and Bacon, 1989; Bacon et al., 1992). Mineral assemblages and compositions are strikingly similar, though Mazama samples contain low amounts of hornblende. Predominance of low crystallinity homogenous rhyodacite is characteristic to both eruptions. Our estimate of pre-eruptive dissolved Hz0 (4.0 +_0.5 wt.%) from Krakatau glass inclusions is virtually identical to estimates from glass inclusions from the Mount Mazama (Bacon et al., 19921, Taupo (Hervig et al., 1989; Dunbar and Kyle, 1993) and rhyolitic portion of the Valley of Ten Thousand Smokes (VTTS) eruptions (Hildreth, 1983, 1987; Westrich et al., 1991). 15.2. Eruption triggering and volatile degassing Injection of basaltic magma at depth in the zoned magma system beneath Krakatau may have triggered the 1883 eruption, as proposed by Self (1992) based on the localized deposition of high alumina basalt ash at the foot of the Perboewatan cone during May 20-26 (Verbeek, 1885; Stehn, 1929). It is unlikely, however, that the basaltic magma would have mixed extensively with rhyodacite because the viscosity contrast between them precludes extensive mixing (Sparks and Marshall, 1986). Alternatively, the replenishment of basaltic parent magma at depth and subsequent convective fractionation (Sparks et al., 1984) may have transferred volatiles (CO, and H,O) and differentiated melt to a growing accumulation of evolved andesite and rhyodacite higher in the Krakatau magma system. Certainly a heat source is required to account for reverse zoned pyroxenes and plagioclase phenocrysts and to maintain low amounts of crystallization in both the gray dacite and upper rhyodacitic magma. Plinian pumice fall deposits marking the onset of the August 26-27 eruption most likely represent tapping of the uppermost, evolved portion of the chamber as gray dacite pumices are rare to absent in these units. Mass discharge rate during the pyroclastic flow phase of the eruption most likely was > 10’

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kg/s. The appearance of gray dacite and banded pumices in the lower pyroclastic flow deposits may be due to increased proportions of more mafic magma in the conduit as a result of increasing discharge rate (Freundt and Tait, 1987) during the flow phase. It is also possible that incipient caldera-forming faults may have transected the sides of the chamber, thereby tapping less-evolved compositions at lower levels in the chamber. Olive obsidians with the highest concentrations of incompatible elements, found within pyroclastic flow deposits may indicate that evolved rhyodaciticrhyolitic magma was present in the uppermost portion of the 1883 chamber. Evidence that this evolved magma may have degassed through conduit walls during earlier May to August eruptive activity (thus sealing off the chamber until the catastrophic eruption) is the 4.0 k 1.0 wt.% volatiles (by difference) found in glass inclusions in obsidian hosts (Table 4). Erosion of the conduit during the flow phase of the eruption may have brought significant amounts of roof and/or conduit lining obsidian to the surface. Strongly banded pumices indicate mingling of two separate magmas in the conduit during eruption. Disequilibrium phenocrysts in white rhyodacite and gray dacite (with highly variable matrix glass compositions on the scale of a hand specimen). and isolated patches of gray dacite in white rhyodacitic hosts indicate more extensive mixing in the chamber probably related to convective fractionation of andesitic parental magma. The petrologic estimate of sulfur discharge from this study is an order of magnitude less than estimates derived from acidity layers in Greenland and Antarctic ice cores (Table 9; Hammer et al., 1980; Legrand and Delmas, 1987; Zielinski, 1995). Even if the contribution from seawater evaporation is considered, the total S discharge estimate of 6.6 X 1012 g S is about half the value reported by Legrand and Delmas (1987), who have corrected for potential seawater sulfate contributions, and slightly over a third of the estimate of Hammer et al. (1980) and Zielinski (1995). A potential source of additional sulfur during the Krakatau eruption is degassing from nonerupted andesitic magma by volatile transfer (H,O, CO, and SO,) to the rhyodacitic portion of the magma chamber perhaps leading to accumulation of an excess vapor phase in the uppermost

C19%)243-274

portion (Sigurdsson et al., 1990a; Gerlach et al., in press). Estimates of total H,O, Cl and S degassing from Krakatau (12.5 km3 DRE, Table 9) are remarkably close to those for the andesitic to rhyolitic VTTS eruption (11.7-15.7 km3 DRE, 6 X 1014 g H,O, 7.8 x 10” g Cl, 1.0 X 10” g S, see Hildreth, 1987; Westrich et al., 1991; Fierstein and Hildreth, 1992). Minor differences in these estimates can be attributed mainly to lower crystallinity and lower relative proportions of andesite and gray dacite in Krakatau erupted products.

16. Conclusions Products from the 1883 eruption of Krakatau total approximately 12.5 km3 (DRE) of magma, 90% of which is composed of white rhyodacitic magma, 4% gray dacite, and 1% andesite (allowing for 5% cognate and accidental lithics). Glass compositions in 1883 erupted products range from andesite (61.6% SiO,) to rhyolite (74.4% SiO,) present as inclusions and in the matrix. Evidence for magma mixing includes banded pumices, glass inclusions that are more mafic than whole-rock and matrix glass compositions, heterogeneous matrix glass compositions on the scale of thin sections, and disequilibrium phenocrysts of plagioclase within both rhyodacite and gray dacite. Incompatible-element enrichments in obsidians and lower rhyodacitic pumice fall samples suggest that the 1883 chamber was compositionally zoned and stratified due to density differences between evolved rhyodacite and more mafic gray dacite or andesite. Gray dacitic pumices represent hybrid compositions resulting from mixing of minor amounts of andesite with larger amounts of evolved rhyodacitic to rhyolitic liquids. The complete trend in major-element chemistry of glass inclusions and matrix glasses can be modelled by 4 1% fractional crystallization of andesitic parental magma to produce the evolved rhyodacitic magma. Plagioclase-liquid equilibria and Fe-Ti oxide compositions indicate that the top portion of the magma chamber was at 1.0 to 1.5 kbar or 4 to 5 km depth, a temperature of 880 to 890°C and log fo, of - 10.7 to - 11 .O. Pre-eruptive dissolved water con-

C. W. Mandeuille et al. / Journal of Volcanology and Geothemal Research 74 (1996) 243-274

centration in 1883 rhyodacite magma was probably 4.0 & 0.5 wt.%, and the 1883 magma was probably not volatile saturated until at a very shallow crustal level of 4-5 km. A petrologic estimate of minimum sulfur discharge from erupted magma is 2.8 X lo’* g S, and minimum chlorine discharge is estimated to be 9.7 X lo’* g Cl. Two potential sources of sulfur from this eruption previously not accounted for may be vaporization of seawater during entrance of pyroclastic flows into the sea and degassing of nonerupted andesitic parental magma lying beneath evolved rhyodacitic magma in the zoned 1883 chamber. Acknowledgements This research was funded by grants from the National Science Foundation (grant # EAR9102229), the National Geographic Society’s Committee for Research and Exploration (grant # NGS-4115-90) and NASA (grant # NAG-51844). We gratefully acknowledge the Indonesian Institute of Sciences (LIPI) for permission to work in Indonesia, and Dr. Sutikno Bronto of the Volcanological Survey of Indonesia (VSI) for providing technical assistance to the project. We thank Charlie Bacon, Jim Luhr and Bruce Marsh for their helpful comments and prompt reviews which greatly improved the manuscript. Mac Rutherford is gratefully acknowledged for financial and technical support during preparation of the manuscript. We thank Rick Murray, Sandi Darter, Don Hermes, Joe Devine, Genji Saito and Noboru Imai for their technical support. References Andersen, D.J. and Lindsley, D.H., 1988. Internally consistent solution models for Fe-Mg-Mn-Ti oxides: Fe-Ti oxides. Am. Mineral., 73: 714-726. Anderson, A.T., 1974. Chlorine, sulfur, and water in magmas and oceans. Geol. Sot. Am. Bull., 85: 1485-1494. Anderson, A.T., 1976. Magma mixing: petrological process and volcanological tool. J. Volcanol. Geotberm. Res., 1: 3-33. Anderson, A.T., Newman, S., Williams, S.N., Druitt, T.H., Skirius, C. and Stolper, E., 1989. H,O, CO,, Cl, and gas in Plinian and ash-flow Bishop rhyolim. Geology, 17: 221-225. Arth, J.G., 1976. Behavior of trace elements during magmatic processes-a summary of theoretical models and their applications. J. Res. U.S. Geol. Surv., 4: 41-47.

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