Journal of Petrology Advance Access originally published online on June 22, 2007
Journal of Petrology 2007 48(8):1543-1567; doi:10.1093/petrology/egm029
© The Author 2007. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: journals.permissions@oxfordjournals.org
Remobilization of Highly Crystalline Felsic Magma by Injection of Mafic Magma: Constraints from the Middle Sixth Century Eruption at Haruna Volcano, Honshu, Japan
Yuki Suzuki1,* and
Setsuya Nakada2
1Institute of Mineralogy, Petrology and Economic Geology, Graduate School of Science, Tohoku University, Aoba-Ku, Sendai, 980-8578, Japan
2Earthquake Research Institute, University of Tokyo, 1-1-1, Yayoi, Bunkyo-Ku, Tokyo, 113-0032, Japan
RECEIVED
JULY 21, 2006;
ACCEPTED
MAY 15, 2007
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ABSTRACT
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The latest eruption of Haruna volcano at Futatsudake took place
in the middle of the sixth century, starting with a Plinian
fall, followed by pyroclastic flows, and ending with lava dome
formation. Gray pumices found in the first Plinian phase (lower
fall) and the dome lavas are the products of mixing between
felsic (andesitic) magma having 50 vol. % phenocrysts and mafic
magma. The mafic magma was aphyric in the initial phase, whereas
it was relatively phyric during the final phase. The aphyric
magma is chemically equivalent to the melt part of the phyric
mafic magma and probably resulted from the separation of phenocrysts
at their storage depth of
15 km. The major part of the felsic
magma erupted as white pumice, without mixing and heating prior
to the eruption, after the mixed magma (gray pumice) and heated
felsic magma (white pumice) of the lower fall deposit. Although
the mafic magma was injected into the felsic magma reservoir
(at
7 km depth), part of the product (lower fall ejecta) preceded
eruption of the felsic reservoir magma, as a consequence of
upward dragging by the convecting reservoir of felsic magma.
The mafic magma injection made the nearly rigid felsic magma
erupt, letting low-viscosity mixed and heated magmas open the
conduit and vent. Indeed the lower fall white pumices preserve
a record of syneruptive slow ascent of magma to 2 km depth,
probably associated with conduit formation.
KEY WORDS: high-crystallinity felsic magma; magma plumbing system; multistage magma mixing; upward dragging of injected magma; vent opening by low-viscosity magma
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INTRODUCTION
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Knowledge of the triggering mechanisms of eruptions is important
for the forecasting of future eruptions. The injection of new
magma into a magma reservoir is one of the major mechanisms.
Theoretical aspects of the triggering process include the increase
in pressure as a result of (1) an increase in the magma reservoir
resulting from the simple volumetric addition of injected magma
(Blake, 1984
), and (2) the intrinsic increase in volume owing
to the exsolution of volatile phases from the melt. Exsolution
can be generated either by cooling-induced crystallization of
the injected high-temperature magma (Folch & Martí,
1998
) or by heating of the low-temperature reservoir magma (Sparks
et al., 1977
). Furthermore, petrological studies on the ejecta
have focused on magma mixing that results from the injection.
To link the magma injection and eruption, some recent petrological
studies have investigated the timescale from mixing to ejection
(e.g. Nakamura, 1995
; Venezky & Rutherford, 1999
). Also,
petrological studies have proposed a new triggering process
for the case of highly crystalline felsic magma eruption: the
injected low-viscosity mafic magma remobilizes the highly viscous
felsic magma, reducing the magma viscosity as a result of mixing
(e.g. Pallister
et al., 1996
; Venezky & Rutherford, 1997
;
Murphy
et al., 2000
; Takeuchi & Nakamura, 2001
).
Although petrological study of ejecta is a powerful tool for studying the triggering mechanisms of eruptions, it has certain drawbacks. There are only a few studies that consider the migration of magma around the reservoir (e.g. Venezky & Rutherford, 1997; Cottrell et al., 1999). Although studies of groundmass microlites can provide information about magma ascent at shallow levels in the conduit (e.g. Hammer et al., 1999; Nakada & Motomura, 1999; Suzuki et al., 2007), they do not address migration around the reservoir. When an eruption is triggered by magma injection, the migration of magma around the reservoir is associated with chemical modification of the magmas, and is thus important for understanding the whole picture of the eruption triggering processes.
The most important factors that characterize and control the migration of the magma are the depth, shape and dimension of the magma storage system. However, it is sometimes difficult to resolve these storage systems, because of the still poor spatial resolution of geophysical imaging methods and incomplete knowledge of the end-member magmas that are mixed (especially the mafic magma, which comes from a deeper level than the felsic magma). The poor resolution of the magma storage systems makes it difficult to evaluate magma migration around the reservoir. To ascertain the locations of magma storage we need to determine the bulk and melt compositions of the end-member magmas, for example, using mass balance of phenocrysts for the mixing of phyric and aphyric magmas (Nakamura, 1995), and the compositions of melt inclusions (Cottrell et al., 1999). If magma storage conditions can be better ascertained, this would help in understanding the interaction of different magmas at the time of injection, through the estimation of their physical properties. At the same time, detailed study of phenocryst zoning could help us to reveal the progress of magma migration and mixing. If no modern observation record exists for a volcano, eruption triggering processes and the depths of magma storage reservoirs, which are obtained from petrological approaches, are particularly important in predicting the future eruptions based on geophysical observations.
With this motivation and background, we have investigated the mid sixth century Futatsudake eruption of the Haruna volcano in central Honshu (Fig. 1). Although no historical documentation is available, the eruption sequence is well known (Oshima, 1983; Soda, 1989, 1996). In this study we show that a highly crystalline (50 vol. %) felsic (andesitic) magma and two mafic magmas (aphyric and phyric) were involved in the eruption. Although the mafic magmas did not erupt without mixing, we reveal that the aphyric mafic magma is chemically equivalent to the melt fraction of the phyric mafic magma and that the two magmas are from a related storage system. We propose that injection of the mafic magma into the felsic reservoir triggered the eruption by forming low-viscosity magmas, such as heated felsic magma and mixed magmas. Based on separate periods of mafic magma ejection (in the first and last phases of the eruption), we conclude that upward dragging of the injected magma can come to an end (see Snyder & Tait, 1996). With regard to magma migration around the reservoir, it is shown that the first erupted magma was associated with conduit formation and ascended slowly. We propose further, using data from detailed compositional zoning in phenocrysts, a multistage mixing model for the formation of homogeneous mixed magma with a minor contribution from mafic magma.
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BACKGROUND OF FUTATSUDAKE VOLCANISM
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Haruna volcano is located in the southern end of the northeast
Japan arc (
Fig. 1). The activity of this volcano (>300 ka)
can be divided into five stages (Oshima, 1983
, 1986
). The ejecta
is andesitic during Stages 1 and 2, andesitic–dacitic
in Stage 3, and dacitic during Stages 4 and 5. Based on the
SiO
2–FeO * /MgO diagram (Miyashiro, 1974
), Watanabe &
Takahashi (1995
) indicated that rocks from Stage 1 activity
can be classified as belonging to the tholeiitic series whereas
those from Stages 2 to 5 belong to the calc-alkali series. During
the first two stages, the formation and destruction of the stratovolcano
were repeated around the present summit (
Fig. 1), forming the
base of the present volcanic edifice. During Stage 3 magma intruded
into the flank of the stratovolcano. Stage 4 is characterized
by two caldera-forming eruptions (40 500 ± 3500 years
BP for the most recent; Oshima, 1986
) at the present summit.
During Stage 5, more than five lava domes (< 0·1 km
3)
were formed inside the caldera and on the eastern flank.
The latest activity occurred at the place where the present Futatsudake lava domes are located (Fig. 1). Three eruptive phases have been identified archaeologically, during the fifth century and in the early and middle sixth century (Machida et al., 1984; Sakaguchi, 1986, 1993). Soda (1989) proposed that the early sixth century event started with a phreatomagmatic eruption, followed by the eruption of pyroclastic flows (Fig. 1; 0·1 km3 in total). The middle sixth century event started with a Plinian eruption, followed by eruption of pyroclastic flows, and ended with lava dome formation. The pyroclastic fall and flow deposits (without any erosion hiatus) are directly covered by stratified fine ash layers from phreatomagmatic eruptions at the time of lava dome emplacement (Soda, 1993, 1996). This indicates that the time elapsed between the first two phases and lava dome emplacement was short. The volumes of the fall and flow deposits for the middle sixth century event are 1·3 km3 and 0·3 km3, respectively (Soda, 1996).
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SAMPLES
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Localities
Pumice blocks of the Plinian and pyroclastic flow phases were
sampled at Owazawa and Ohinata (
Figs 1 and
2). The samples were
classified as ‘white pumice’, ‘gray pumice’
and ‘banded pumice’ (a mixture of the magmas forming
the white and gray pumices). In Owazawa, the pyroclastic fall
deposit was divided into the lower, middle and upper fall deposits,
each of which is distinctive in terms of the type and size of
the pumice. The pyroclastic flow deposit has been divided into
the lower flow, top 30–60 cm and top 0–30 cm. The
lower fall unit includes gray and banded pumices along with
white pumices. In Ohinata, the pyroclastic fall deposit has
been divided into five units (
Fig. 2). The gray and banded pumices
are limited to the lowermost unit (unit 1;
Fig. 2), similar
to Owazawa. The samples from Ohinata were used only for bulk-rock
analyses. The lava dome samples were collected from three peaks
(Odake, Medake and Magodake).
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Fig. 2. Columnar sections showing eruptive sequence and bulk-rock SiO2 contents of juvenile materials. Distance from Futatsudake is shown next to the site name. For the Owazawa site, SiO2 contents for white pumices are shown with different symbols, depending on the number of pumices used for a single analysis.
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Petrography
Crystals in pumice and dome lava include phenocryst phases and
microlites in the groundmass. The phenocryst phase can be classified
as ‘phenocryst’ (>300 µm across) and ‘microphenocryst’
(100–300 µm across) (
Fig. 3). The size difference
between microphenocrysts and microlites is evident only in the
white pumices and white parts of banded pumice (microlites of
<20 µm across). Amphibole exists, but not as a microlite
(
Table 1).
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Fig. 3. Photomicrographs of representative ejecta from the middle sixth century eruption at Futatsudake (plane-polarized light). All scale bars represent 1 mm. (a) Banded pumice from the lower fall deposit; (b) white pumice from the lower fall deposit; (c) white pumice from flow top 30–60 cm; (d) dome lava (Odake). In (a), the dark gray part (G) is adjacent to the white part (W), and the boundary between the two is sharp. Cpx, clinopyroxene; Am, amphibole; Opx, orthopyroxene; Pl, plagioclase; Ox, Fe–Ti oxide. (Note clinopyroxene with different sizes from microphenocrysts (MPh) in (b) and phenocrysts in (d).)
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The common phenocryst phases for all ejecta are plagioclase
(< 2 mm), orthopyroxene (< 2 mm), amphibole (< 5 mm),
magnetite (< 1 mm) and ilmenite (< 700 µm) (
Table 1;
Fig. 3). Sometimes they form multi-phase aggregates (
Fig. 3c),
and contain inclusions of other minerals and glass. Olivine
occurs as a phenocryst (< 2 mm) and microphenocryst in the
dome lava (
Table 1;
Fig. 4), and is found as a microphenocryst
(
100 µm across) in the gray pumice and the gray part of
the banded pumice (
Table 1;
Fig. 4). Clinopyroxene is found
as a phenocryst (<500 µm;
Fig. 3d) and as a microphenocryst
in the dome lava, and as a microphenocryst (< 300 µm;
Fig. 3b) in the white pumice and the white part of the banded
pumice in the lower fall deposit (
Table 1). Olivine and clinopyroxene
contain inclusions of Fe–Ti oxides.
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Fig. 4. Backscattered electron image of typical olivine (Ol) in ejecta from the middle sixth century eruption at Futatsudake. Scale bars represent 100 µm. (a) Phenocryst in dome lava (Odake); (b) microphenocryst olivine in gray pumice from the lower fall. (Note the size difference between (a) and (b), and euhedral outline of the crystal in (a) and (b) irrespective of the orthopyroxene (Opx) reaction rim (R.R.).)
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Although phenocrysts and microphenocrysts are commonly euhedral
and free from reaction rims, this is not the case for the following
phases (
Table 1). Ilmenite shows resorption in gray pumice and
the gray part of banded pumice. Olivine has an orthopyroxene
reaction rim (
Fig. 4). Some plagioclase phenocrysts and microphenocrysts
in the dome lavas and the lower fall (white, gray and banded)
pumices have dusty zones (
100 µm wide). Some orthopyroxene
grains in the dome lava, gray pumice and gray part of banded
pumice have clinopyroxene reaction rims (
20 µm). Some
of the grains of amphibole are mantled by aggregations of pyroxene,
plagioclase, Fe–Ti oxide and glass (up to
100 µm
wide), as seen in pumices from the lower fall, upper fall and
flow top 30–60 cm. All amphibole grains are mantled by
similar rim intergrowths in the dome lava.
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ANALYTICAL METHODS
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Major and trace elements in minerals and glass were analyzed
with a JXA-8800R EPMA at the Earthquake Research Institute (ERI),
University of Tokyo. The beam diameter was set focused for minerals
and at 10 µm for glass. Major elements were analyzed using
an accelerating voltage of 15 kV and a beam current of 12 nA.
The measuring time was 10 s for the peak and the two backgrounds,
respectively, for every element. In glass analyses, Na, K and
Si were measured in the first analytical cycle and Al in the
second cycle, to minimize the evaporation of alkalis as a result
of continuous beam exposure (Devine
et al., 1995
). For NiO in
olivine, the accelerating voltage and counting time were set
at 20 kV and 120 s, respectively. Microlites with a large enough
size were analyzed by EPMA. At least, two samples of each type
from the individual eruptive units were analyzed.
Water contents in melt inclusions in phenocrysts were determined using the EPMA difference method (Devine et al., 1995; Morgan & London, 1996), using the JXM-8800 EPMA in Tohoku University. Hydrous glass standards were used as reference, together with the unknown. Following the optimal conditions for hydrous glass analyses (Morgan & London, 1996), the beam diameter was set at 15 µm, the beam current at 2 nA and the measuring time at 30 s (for both the peak and the two backgrounds). The rest of the analytical conditions were similar to those used for the mineral and glass analyses. Vesiculated inclusions were not used for water content estimation.
Whole-rock compositions were analyzed by X-ray fluorescence (XRF) (PW2400) at ERI, using glass beads with 10 parts flux to one part sample. When pumice grains are not large enough for X-ray analysis, multiple grains having similar characteristics and from the same horizon were used (Fig. 2). We could not separate each band of the banded pumices and hence obtained their bulk compositions (Table 2). Four to seven dome lava samples were analyzed for each peak of Futatsudake.
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Table 2: Representative whole-rock analyses for pumice and dome lava samples. All oxide values normalized to 100% and total iron as FeO
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ANALYTICAL RESULTS
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Whole-rock chemistry
The SiO
2 content of the white pumice is nearly constant, independent
of the eruptive unit (60·0–61· 6;
Fig. 2).
Gray and banded pumices have lower SiO
2 content (57·5–60·4
wt %) than the white pumice. The SiO
2 content of the banded
pumice (57·5–60·4 wt %) is similar to that
of the gray pumice (57·9–58·5 wt %) or higher.
The SiO
2 content of the dome lava (59·7–60·4
wt %) is lower than that of the white pumice, but partly overlaps
with that of the banded pumice (
Fig. 5). All samples show a
linear trend in most Harker diagrams (
Fig. 5). In other diagrams
(e.g. MgO–SiO
2, K
2O–Cr), however, the trend for
gray and banded pumices is different from that of the dome lava.
Mineral chemistry
Compositions of the major phenocryst phases are reported in
Electronic Appendices 1–6, available as Supplementary
Data at
http://www.petrology.oxfordjournals.org. In the following
discussion, ‘Rim’ represents the outermost part
of a crystal adjacent to the groundmass or reaction rim. The
part inside the rim is referred to here as the ‘core’,
but we introduce different terms for orthopyroxene and dusty
plagioclase that shows complex zoning. Cores of olivine, clinopyroxene,
magnetite and ilmenite are homogeneous except near the rims,
so the core data are from crystal centers in the thin sections.
Olivine
Olivine crystals all show normal zoning in Mg-number [100 x Mg/(Mg + Fe)] (Fig. 6a). The olivine microphenocrysts in the gray pumice and gray parts of banded pumice of the lower fall deposit have a core composition lower in Mg-number (78–76) than the phenocrysts and microphenocrysts in the dome lava (Fig. 6a). In the Mg-number vs NiO diagram, the composition of olivine in dome lava overlaps with that of the olivine microphenocrysts from the lower fall deposit (Fig. 6b).
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Fig. 6. Chemistry of olivine in gray pumice, gray part of banded pumice (BP), and dome lava. (a) Mg-number [Mg/(Mg + Fe)] of cores vs Mg-number of rims; (b) Mg-number vs NiO. Legend in (a) also applies to (b). Data for (b) were obtained through line-scan analyses for representative grains.
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Clinopyroxene
Phenocrysts and microphenocrysts in the dome lava have cores
of Mg-number 80–75 and Wo [Ca/(Mg + Fe + Mn + Ca)] 46–40,
and rims of Mg-number 74–70 and Wo 45–38, showing
normal zoning (
Fig. 7). Clinopyroxene in the overgrowth rim
on orthopyroxene phenocrysts in gray pumices and gray parts
of banded pumices has Mg-number 71–74 and Wo 35–41.
Orthopyroxene
Two types of orthopyroxene can be identified based on the zoning
profile (
Fig. 8). Type 1 has an Mg-number peak in the rim, whereas
Type 2 has the peak inside the rim. Regardless of the type,
each phenocryst has a homogeneous part (‘core’)
accounting for most of the area. The cores of both types are
almost identical in composition. For Type 2, we define the part
between the ‘core’ and the ‘rim’ as
‘inner rim’. Occurrence of Type 2 is limited to
the lower fall pumice and the dome lava (
Figs 8 and
9). The
core composition range (Mg-number 62·4–65·7
and Wo 0·8–1· 8) is uniform throughout the
eruption. Rims of Type 1 have Mg-number 65·3–67·5
and Mg-number 65·3–67·8 for the gray and
white parts of pumices (including banded pumices), respectively,
for lower fall, and Mg-number 66·0–73·2
for the dome lava. These values are higher than those for others
(Mg-number 63·8–66·0). Inner rims of Type
2 have higher Mg-number than the rims of Type 1 in each eruption
phase. Orthopyroxene microlites in the groundmass overlap chemically
with the phenocryst rims (
Fig. 9).
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Fig. 8. Results of line-scan profiles across representative orthopyroxene phenocrysts. (For explanation of phenocryst types, see the text.) ‘Cpx’ indicates reaction rim (no data shown). Although the core centers are not shown in these profiles, the compositions are almost uniform throughout the core parts.
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Fig. 9. Mg-number vs Wo [Ca/(Mg + Fe + Mn + Ca)] mol% plot for orthopyroxene phenocryst (Ph) and groundmass microlites. (For explanation of types and definition of parts in phenocryst, see the text.) BP, banded pumice.
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Amphibole
Most amphibole phenocrysts are ‘hornblende’ (
Fig. 10).
They have oscillatory zoning in Na + K (in A site) and Al(IV),
both of which change in a sub-parallel fashion. The rims are
relatively low in Na + K (
Fig. 10) and Al(IV). No systematic
chemical changes were found between eruption units.
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Fig. 10. Si vs Na + K (in A site) diagrams for amphiboles in representative samples. Dotted line indicates the boundary between tschermakite (Ts) and hornblende (sénsu strícto) (Hb), after Deer et al. (1992). Data for dome lavas do not include those of oxyhornblende. We analysed crystals without breakdown rims (Table 1) for samples excluding the dome lava. Rim composition is not available for the dome lava because of breakdown (Table 1). The compositional range in the middle fall deposit and lower flow is slightly wider than in the lower fall and is probably due to the larger size of the phenocrysts. There are no large phenocrysts in the lower fall deposit because of fragmentation.
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Plagioclase
For a dusty phenocryst, we define parts inside the dusty zone
as ‘core’, and the parts between ‘core’
and the ‘rim’ as ‘inner rim’ (
Figs 11 and
12). Cores of all the plagioclase phenocrysts show oscillatory
zoning, ranging from 55 to 90 in An mol % (100
x Ca/(Ca + Na)]
(
Figs 11 and
12). The FeO and MgO contents are relatively homogeneous
with 0·2–0·4 wt % and
0·05 wt %,
respectively, except near the rims of some clear phenocrysts
in lavas and lower fall pumices. Analytical uncertainties (relative
percent) in FeO and MgO are
c. 20% at 0·5 wt % and
c.
40% at 0·05 wt %, respectively. The rims of clear phenocrysts
vary in their chemistry within the range of their cores, except
for the lower fall pumice and dome lava. Cores of both clear
and dusty plagioclase phenocrysts in the lower fall pumice and
dome lava mostly have lower FeO and MgO contents than in the
rim and inner rim (
Figs 11a, b and
12). Plagioclase microlites
in the groundmass overlap chemically with the rim and inner
rim of plagioclase phenocrysts (
Fig. 12).
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Fig. 11. Line profiles for representative plagioclase phenocrysts. Each line profile is shown by an arrow in the adjacent backscattered electron image. Contents of An, FeO and MgO are shown on the left and right vertical axes, respectively. Shaded area in (a) indicates the inner rim and rim of dusty plagioclase. Although the core centers are not shown in these profiles, the compositions are almost uniform throughout the core parts. (For explanation of phenocryst types (clear, dusty) and definition of parts in phenocrysts, see the text.) Pl, plagioclase; DZ, dusty zone.
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Fig. 12. An (mol %) vs FeO and MgO (wt %) diagrams for plagioclase phases. FA, fall; LFA, lower fall; GP, gray pumice; WP, white pumice; BP, banded pumice; G, gray part; W, white part; FL, flow; DL, dome lava; Ph, phenocryst. (For explanation of phenocryst types (clear, dusty) and definition of parts in phenocrysts, see the text.)
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The thickness of the final stage of growth is commonly larger
in microphenocrysts than in phenocrysts because of the different
surface areas. The wider domain makes it possible to determine
precisely, by microprobe analysis, the compositions of the parts
equilibrated just before the eruption. The An contents of microphenocryst
rims are compared in
Fig. 13. Microprobe backscattered electron
images show that the microphenocryst rims are different chemically
from the groundmass microlites in these samples, even if microlites
exist. Distribution of An content does not change significantly
from sample to sample. FAL-2 and FAL-5 from the lower fall deposit
have the lowest An contents for both maximum and minimum values
of the whole distribution (An
50 and An
70, respectively).
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Fig. 13. An (mol %) histograms for plagioclase microphenocryst rims. The data in each histogram are from a single sample of white pumice. Samples with asterisks include plagioclase microlites in the groundmass. The dimensions of the vertical axes are the same in all the histograms. FA, fall; LFA, lower fall; FL, flow. However, the microprobe backscattering images show that the phenocryst rims are different chemically from groundmass microlites.
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Magnetite
The cores of magnetite phenocrysts are chemically homogeneous,
with x-usp 0·17–0·20 (where x-usp is the
fraction of the ulvöspinel molecule) and Mg/Mn 3–8,
except for some grains in the lower fall gray parts of the banded
pumices with high x-usp (
Fig. 14). Phenocryst rims show no difference
in x-usp and Mg/Mn from their cores, except for the lower fall
pumices and lavas. In the white parts of pumices (including
banded pumices) of the lower fall and dome lavas, some of the
phenocryst rims have higher x-usp, but have similar Mg/Mn, in
comparison with the cores. In the gray pumice and gray part
of the banded pumice, the rims of some phenocrysts have higher
x-usp and Mg/Mn. The cores of the microphenocrysts seem to have
crystallized simultaneously with the outer zones of the phenocrysts.
For example, zoned margins (rim and near rim) of phenocrysts
chemically resemble the cores of the microphenocrysts.
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Fig. 14. Histograms for x-usp and Mg/Mn contents of magnetite phenocrysts (Ph) and microphenocrysts (MPh). C, core; R, rim; FA, fall; LFA, lower fall; GP, gray pumice; WP, white pumice; BP, banded pumice; G, gray part; W, white part; FL, flow; N, number. (Note the different x-usp axis values for the dome lava.)
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Ilmenite
Cores of ilmenite phenocrysts are chemically homogeneous, with
x-ilm 0·55–0·60 (where x-ilm is the fraction
of ilmentite in the rhombohedral phase) and Mg/Mn 6–9
(
Fig. 15). The rims of some phenocrysts in the lower fall unit
(white, gray and banded pumices) have higher x-ilm and Mg/Mn
(
11) than the cores. However, such an increase is not found
in the other eruptive phases. The cores of the microphenocrysts
are similar in composition to the phenocryst cores.
Fe–Ti oxide thermometry and fO2 barometry
Because ilmenite is resorbed (Table 1), the ilmenite rims in the gray part of the pumice (including banded pumice) were not used. The equilibrium between the ilm-mt pairs used for calculation was checked in terms of Mg/Mn (Bacon & Hirshmann, 1988). The algorithm of Andersen & Lindsley (1988) yields temperatures of 820–850°C and log fO2 of –10·3 to –10·9 for the core pairs from all samples and the rim pairs without compositional changes from the cores. The algorithm yields a temperature of 880°C and log fO2 of –9·9 to –10·8 for the rim pairs from the white pumice fraction (including banded pumice) of the lower fall deposit.
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Fig. 15. Range of x-ilm compositions for ilmenite phenocrysts (Ph) and microphenocrysts (MPh). Data are not available for samples including dome lavas because of their exsolution (Table 1) and scarce amount of crystals. FA, fall; FL, flow. White part of pumice of the lower fall includes white pumice and white part of banded pumice. Similarly, gray part of pumice includes gray pumice and gray part of banded pumice. Data for banded pumice are connected with dotted lines.
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Groundmass composition
We employed two methods to obtain the composition of the groundmass.
For crystal-free groundmass samples (white pumice;
Table 1)
with relatively large glass domains (
10 µm across) the
analysis of the glass part was treated as the groundmass composition.
For crystal-bearing groundmass (
Table 1), we averaged analytical
data for up to 200 point analyses. The latter method was limited
to lavas (three representatives), as the highly vesiculated
groundmass of pumice prevents random analyses. The estimated
SiO
2 content of the groundmass for the white pumice is in the
range of 77·4–78·4 wt % (
Table 3); no systematic
change between eruption units was found. The groundmass of the
dome lava is less fractionated (SiO
2 of 72–73 wt %) than
that of the white pumice.
Phenocryst phase volume
For highly vesiculated samples it was necessary to section many
samples to obtain statistically reasonable estimates of the
phenocryst mode. Accordingly, we employed chemical mass balance
only for the white pumice (with high vesicularity) in all samples.
The middle fall white pumice we used as representative (
Table 5)
because it is free from groundmass microlites, and its groundmass
glass is considered equivalent to the melt part of the reservoir
magma. The total volume of the phenocrysts (including microphenocrysts)
decreases from the white pumice (
c. 50%), to the dome lava (37·0–48·8%)
and to the gray pumice (30·5–35·4%) (
Tables 4 and
5). Ratios of individual phenocrysts to total phenocrysts
barely change throughout the samples (e.g. plagioclase is consistently
the most abundant phase). Olivine and clinopyroxene account
for
0·6 vol. % in total for the dome lava (
Table 4).
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DISCUSSION
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Mixed and heated magmas at the beginning and end of the eruption
Constraints from mineralogy
The lower fall pumice and the dome lava have characteristics
that cannot be explained by simple equilibrium crystallization
from one to the other. It is important to note here that the
gray and white parts of the banded pumice are identical to the
white pumice and gray pumice of the lower fall deposit. We stress
the following properties:
- coexistence of normally and reversely zoned mafic phases (olivine, clinopyroxene, and orthopyroxene) in the gray parts of the pumice (including banded pumice) and the dome lava (Figs 6–9
);
- reversely zoned orthopyroxenes in the white parts of the pumice (including the banded pumice) (Figs 8 and 9);
- chemical disequilibrium between olivine and orthopyroxene in the gray parts of the pumice (including the banded pumice) and the dome lava (Fig. 16a);
- enrichment in FeO and MgO at and near the rims of plagioclase (Figs 11a, b and 12);
- Mg/Mn increasing at and near the rims of both magnetite (gray parts of pumice including banded pumice; Fig. 14) and ilmenite (gray and white parts of pumice including banded pumice).
Compositional zoning, as in (1), (2), (4) and (5), was formed
just before eruption, because it is found near the rims. The
characteristics (2), (4) and (5) occur when crystals encounter
less evolved melt. The most plausible process for explaining
these characteristics would be magma mixing. The onset of amphibole
and magnetite precipitation (Gill, 1981
) or an increase in
fO
2 may produce reverse zoning of mafic minerals. However, in such
cases, the lines of evidence (4) and (5) favor magma mixing
instead. FeO in plagioclase can increase as a result of an increase
in
fO
2 (Hattori & Sato, 1996
). However, a simultaneous increase
of MgO does not support a change in
fO
2. Formation of a dusty
zone in plagioclase requires an encounter with more calcic melt,
an increase in temperature (e.g. Nakamura & Shimakita, 1998
),
or perhaps both, in accordance with the increases in FeO and
MgO. The increase in Mg/Mn supports a temperature increase as
a result of mixing with high-temperature mafic magma. Dissolution
of ilmenite in the gray pumice and the gray part of banded pumice
(
Table 1) could be related to the mixing process. Normally zoned
phenocrysts of olivine and clinopyroxene were derived from the
high-temperature magma, whereas plagioclase and Fe–Ti
oxides are from the low-temperature magma (
Fig. 17). Microphenocryst
olivines in the gray pumice and the gray part of the banded
pumice are not crystals that nucleated upon and after mixing,
because they have reaction rims of orthopyroxene. The fact that
the microphenocryst olivine in the gray part of the pumice (
100
µm) is much smaller in size than olivine in the dome lava
(
Fig. 4) may represent a shorter growth time prior to mixing;
the mafic magma with olivine microphenocrysts (
100 µm)
having been aphyric until just before mixing. Amphibole is from
the low-temperature magma, because of the association with aggregates
of plagioclase and Fe–Ti oxides. However, their rims lack
a record of heating, as indicated by increases in Na + K and
Al(IV) (Blundy & Holland, 1990
; Scaillet & Evans, 1999
;
Rutherford & Devine, 2003
), in the lower fall white pumice
(
Fig. 10). This may be due to the amphibole becoming unstable
and no longer growing after the mixing event (
Table 1). The
above phenocryst assemblages in the end-member magmas are also
supported by chemical equilibrium or phase stability between
the phenocrysts. That is, olivine and clinopyroxene in the dome
lava show equilibrium Mg–Fe distributions (
Fig. 16b).
The white pumice and the white part of the banded pumice from
the lower fall deposit contains clinopyroxene microphenocrysts
(
Table 1) that probably nucleated during or after mixing, because
they have no chemical zoning or textures indicative of mixing.
Even in the gray pumice and the gray part of the banded pumice,
plagioclase was not present in the high-temperature end-member
(
Fig. 17). If so, some plagioclase microphenocryst cores should
have higher MgO and FeO than the plagioclase rims crystallized
from the mixed magma (absent in
Fig. 12).
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Fig. 17. Schematic diagram showing the range of magmas involved in the Futatsudake eruption, together with their phenocryst assemblages. ‘White’ indicates the white part of pumice including banded pumice; ‘Gray’ indicates the gray part of pumice including banded pumice; Ol, olivine; Cpx, clinopyroxene; Opx, orthopyroxene; Am, amphibole; Mt, magnetite; Il, ilmenite; Pl, plagioclase.
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White pumice from the middle fall deposit to the flow top 0–30
cm probably maintains the original characteristics of the low-temperature
end-member magma (
Fig. 17), considering the phenocryst assemblage
and the chemical compositions. Thus, the magma erupted in the
middle stage represents the low-temperature end-member of the
mixed magmas. In the pumice, a slight reverse zoning of orthopyroxene
(
Figs 8 and
9) is observed, whereas enrichments of FeO, MgO
and Mg/Mn in the phenocryst rims of plagioclase and magnetite
are not found (
Figs 11,
12 and
14). The following two scenarios
may explain this. The first is that the reverse zoning of orthopyroxene
resulted from an increase in
fO
2. The loss of H
2 gas resulting
from devolatilization causes oxidation (e.g. Czamanske &
Wones, 1973
), which is likely to occur during syneruptive degassing
processes. However, an increase of
fO
2 should increase the FeO
content of plagioclase (Hattori & Sato, 1996
), which is
not found in the present case (
Figs 11 and
12). The other possibility
may be that the reverse zoning in orthopyroxene records a previous
heating event not shown by the magnetite because of rapid diffusion
re-equilibration. In such a scenario, it is possible that reverse
zoning of orthopyroxene was preserved, whereas plagioclase rims
that once had been enriched in FeO and MgO disappeared because
of diffusion. At the temperature of the felsic magma (
850°C),
the diffusion coefficient of Mg in plagioclase with An
70 [=
1·3
x 10
–18 m
2/s, calculated using the data of
Costa
et al. (2003
)] is found to be larger than the Fe
2+–Mg
interdiffusion coefficient in orthopyroxene (2·0
x 10
–20 m
2/s; Tomiya & Takahashi, 2005
). We shall defer discussion
on the possible heating of the felsic magma erupted in the middle
stages of the eruption until the section ‘Possible pre-heating
of the felsic reservoir magma’.
Constraints from groundmass and bulk-rock compositions
The linear bulk-rock trends for the gray pumices and dome lavas (Fig. 5) support their origin as binary-mixing products. However, the trends for the lavas and gray pumices intersect on the high K2O side of the K2O–Cr diagram, indicating the same low-temperature (high K2O) mixing end-member but different high-temperature end-members (phyric and aphyric mafic magmas; Fig. 17).
Although the white pumice is considered in terms of bulk chemistry to be the felsic end-member of mixing (Fig. 5), the white pumice of the lower fall deposit has some characteristics of a mixed magma; for example, the properties (2), (4) and (5) in the previous section. The white parts of the banded pumices have the same mineralogical characteristics as the lower fall white pumices. For these units, we suggest heating in a chemically closed system instead of mixing (Fig. 17). The heating can make the melt composition less evolved as a result of dissolution of earlier formed crystals, which leads to the formation of the crystal textures and compositional zoning observed in mixed magmas. Neighboring mixed magma (gray pumices and gray parts of banded pumices) or its mafic end-member should have heated the felsic magma. Irrespective of the heating that may result in the partial dissolution of phenocrysts, the white pumices from the lower fall deposits have similar groundmass compositions to the white pumices of the middle eruption (Table 3). This can be explained by the possibility that the lower fall felsic magma experienced slow syneruptive ascent and resultant crystallization at shallow level in the conduit (discussed below).
In major element variation diagrams the banded pumices plot on a mixing line between white and gray pumices (Fig. 5). This is consistent with their mineralogical characteristics; the gray and white parts of the banded pumices resemble the white pumice and gray pumice of the lower fall, respectively, and the resultant banded pumice is equivalent to a mixture of the white and gray pumice.
Heterogeneous reactions of phenocrysts upon heating and mixing in the lower fall deposits and dome lavas
Patterns of compositional zoning and the degree of compositional change are variable in each product, and are even discernible at the hand-specimen scale. Examples are Types 1 and 2 orthopyroxenes (Figs 8 and 9), clear and dusty plagioclases (Figs 11 and 12) and magnetites with variable x-usp and Mg/Mn values (Fig. 14). These indicate heterogeneous physicochemical changes and/or variable times from such changes to the magma quenching. Because of the extent of magma mingling or mixing, this heterogeneity should be anticipated and it might reflect the proximity of the contrasting magmas. The coexistence of variable phenocrysts in a hand specimen implies effective stirring in each mixed and heated magma.
Timescale from mixing and heating to final ejection for the lower fall deposit
We provide a rough estimate of mixing timescales by using the zoning pattern of magnetites in the white pumices and white parts of banded pumices from the lower fall deposit. The zoning profile in the magnetite grains (200 µm across) would be homogenized in 1·5–12 years at 800–900°C (Tomiya & Takahashi, 2005). Magnetite microphenocrysts in the white pumices and in the white parts of the banded pumices from the lower fall deposit have relatively homogeneous cores (x-usp 0·17–0·20; Fig. 14), surrounded by reversely zoned rims that are as thin as <5 µm. The core compositions are similar to those of the white pumices from the middle eruption, whose magma represents the low-temperature end-member magma in the mixing (Fig. 17). The cores are believed to record the condition before the heating of the lower fall magma. The preservation of the core composition represents a timescale from heating (mixing) to eruption that is shorter than 1·5–20 years. This indicates that the heating and mixing are caused by an event unrelated to an eruption that took place in early sixth century at Haruna, 50 years before the eruption studied here. Devine et al. (2003) showed experimentally that magnetites in the Soufrière Hills magma became zoned within 2–10 days after temperature changed from 835–850 to 860–880°C. This condition may provide a minimum timescale for the white pumice and the white part of banded pumice.
Characteristics of the mafic end-member magmas
Chemical and genetic relationships between the mafic magmas
The linear trend in Fo–NiO of olivines (Fig. 6b) suggests that the aphyric and phyric mafic magmas (Fig. 17) have a close genetic relationship. However, the switch from aphyric to phyric magma during an eruption does not simply result from the temporal evolution of aphyric mafic magma as a result of crystallization. If so, small olivine crystals in the gray pumices and gray parts of banded pumices should have the same core composition as the phenocrysts in the dome lavas (not observed in Fig. 6a) and mixed magmas, formed from the common felsic end-member magma (Fig. 17), should have the same bulk-rock compositional trend (not observed in Fig. 5). It is instead considered that the aphyric mafic magma is chemically equivalent to the melt part of the phyric mafic magma. The cores of the small olivine microphenocrysts in the gray pumice and the gray part of the banded pumice are chemically close to the rim of the olivine phenocrysts in the dome lava (Fig. 6a). This indicates that small olivines in the gray pumices and the gray parts of the banded pumices nucleated from a melt of similar composition to that which crystallized the rim of the phenocrysts.
Bulk composition of the aphyric mafic magma
The bulk chemistry of the aphyric mafic magma was estimated using phenocryst volumes and the major oxide compositions of the mixed magma (gray pumice) and felsic magma (white pumice). Assuming 0 vol. % phenocrysts in the aphyric magma, the estimated compositions are shown in Table 6 (A and B; c. 52 wt % SiO2). The chemistry resembles that of basalt erupted during the early stage (Stage 1) of Haruna activity (Table 6), which has chemical characteristics common to this volcano. This estimation is supported by observed equilibrium between the melt (aphyric bulk) and the cores of olivine microphenocrysts in the gray pumice and the gray part of the banded pumice. KD for the Fe–Mg distribution [(Fe/Mg)Ol/(Fe/Mg)melt] is 0·31–0·36 for the two compositions of aphyric mafic magma (A and B in Table 6). We used 0·1 for Fe3+/(Fe2+ + Fe3+) in the aphyric mafic magma at quartz–fayalite–magnetite (QFM) and nickel–nickel oxide (NNO) oxygen buffer conditions, based on estimations from Sack et al. (1980). For equilibrium olivine–melt pairs, the KD can be 0·3 ± 0·03 for a wide range of melt compositions, temperatures and water contents (e.g. Roeder & Emslie, 1970; Ulmer, 1989).
Phenocryst content and bulk composition of the phyric mafic magma
Mass-balance calculations for the formation of the dome lava
through the addition of aphyric magma (melt) and olivine and
clinopyroxene phenocrysts (derived from the phyric mafic magma)
to the white pumice magma yields their mixing ratios (
Table 7).
The ratios of felsic magma : aphyric mafic magma (melt) : olivine
: clinopyroxene are about 85–86 : 12–13 : 0·3
: 0·3 (in wt %). The weight per cent ratios of olivine
and clinopyroxene are roughly consistent with their phenocryst
volumes in the dome lava (
Table 4), when considering their relative
densities. In the calculation, phenocrysts of felsic magma origin
account for 42–44 vol. % of the mixture, consistent with
the total phenocryst volumes in the dome lavas (37–48
vol. %;
Table 4). The calculated proportion of phenocrysts in
the phyric mafic magma is about 6 wt % (
Table 7). These results
allowed us to estimate the bulk composition of the phyric mafic
magma (
Table 6).
Figure 18 shows the relationships between mixed magmas (gray
and banded pumices, and dome lavas) and the mafic end-members
(aphyric and phyric), in which the phyric and aphyric mafic
magmas lie at the ends of the mixing lines. The enrichment of
MgO and Cr content in the phyric mafic magma compared with the
aphyric mafic magma can be explained by the concentration of
olivine and clinopyroxene in the former, in both of which Cr
is highly compatible (Gill, 1981
).
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Fig. 18. Mixed magmas (dome lava, gray pumice) and the mafic end-members of mixing in MgO–SiO2 (wt %) and K2O(wt %) –Cr (ppm) variation diagrams. Each regression line determined for the mixed magmas (gray pumices or dome lavas) and the felsic end-member (white pumice, not shown in this figure) connects mafic and felsic end-member magmas. Phyric mafic magma is compositionally equivalent to a mixture of aphyric mafic magma and phenocrysts (Table 7). (a) A and B are given in Table 6. (b) Mafic magmas are plotted, based on estimated K2O contents (Table 6, average of A and B).
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Storage conditions of the magmas
Storage condition of the mafic magmas
Because olivine and clinopyroxene were stable in the phyric
mafic magma, the stability conditions were estimated using the
MELTS algorithm (Ghiorso & Sack, 1995
), using the composition
of the melt fraction (aphyric mafic magma;
Table 6) with different
water contents. The estimated liquidus surface shown in
Fig. 19 indicates near-liquidus coexistence of olivine and clinopyroxene
at pressures of more than 4 kbar and with 2–6 wt % of
water, varying with pressure. Under these conditions, the compositions
of the crystallized olivine and clinopyroxene (Mg-number of
77–79 and 77–80, respectively) are in the composition
range of the dome lavas (
Figs 6a and
7). The calculated liquidus
temperature (1130–1150°C) shows no correlation with
pressure. The clinopyroxene barometer (Nimis & Ulmer, 1998
)
using phenocryts in the dome lavas yields 3–4 kbar pressure
at 1130°C. This value is in agreement with the requirement
from
Fig. 19. Thus, we consider this value of
4 kbar as the
storage pressure for the mafic magma (
Fig. 20).
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Fig. 19. Liquidus surface of aphyric magma (Table 6, A), estimated using MELTS (Ghiorso & Sack, 1995) as a function of pressure and H2O content. Redox conditions appropriate to QFM were assumed because the oxygen fugacity of typical island arc magma is in the range QFM–NNO + 2 (Carmichael & Ghiorso, 1990). Although a decrease in fO2 expands the field of olivine relative to orthopyroxene, the stabilities of clinopyroxene and plagioclase are independent of fO2 (e.g. Berndt et al., 2005). Therefore, we may have overestimated the olivine stability field if the fO2 is more oxidizing than assumed. Pl, plagioclase; Px, pyroxene; Cpx, clinopyroxene; Opx, orthopyroxene; Ol, olivine. The shaded zone shows olivine–clinopyroxene coexistence on the liquidus.
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Fig. 20. Plausible magmatic processes before and during the Futatsudake eruption. LFA, lower fall; FA, fall; FL, flow. The erupted felsic magma originated from a part of the reservoir that was poorer in crystals and less viscous. The less viscous part might have been formed by pre-heating of the reservoir. After injection of mafic magma from beneath (step 1), mixed and heated magmas were forced to move upward by the convection driven by the injection of the hot mafic magma (step 2). The movement stopped because of cessation of injection and because of cooling (step 3). After establishment of an open conduit, crystal-rich felsic magma was erupted (step 4). Dome lavas sampled at the surface are the products of magma mixing. However, we cannot discount the existence of rocks as felsic as white pumice in the lower part of the lava dome (arrow in step 5). (For possible scenarios of dome lava formation (step 5), see text.)
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The estimated densities of melt, olivine and clinopyroxene under
the above conditions are 2600 kg/m
3, 3400 kg/m
3 and 3300 kg/m
3.
This means that there is a high possibility of density-segregation
of mafic minerals from the melt in the magma reservoir. It is
plausible that aphyric mafic magma was involved during the early
eruption phase (
Figs 17 and
20), provided that it was in the
upper portion of the reservoir. Assuming a homogeneous upper
crust having a density of 2700 kg/m
3(e.g. granite–diorite,
Best & Christiansen, 2001
), a pressure of 4 kbar would correspond
a depth of
c. 15 km (
Fig. 20).
Pressure in the felsic magma reservoir
As the white pumices in all eruptive phases are homogeneous in terms of their phenocryst cores, they are considered to be derived from the same reservoir. We have determined the water content in the reservoir magma by analyzing the water content in melt inclusions in phenocrysts in two representative samples (middle fall and lower flow). The average water content of vesiculation-free melt inclusions in orthopyroxene and plagioclase phenocrysts is 5·3 ± 0·4 wt % (n = 14). This corresponds to the solubility of water in the felsic melt at 1·9 ± 0·3 kbar according to Burnham's model (Holloway & Blank, 1994). It is therefore reasonable to assume that the felsic magma was water-saturated, based on evidence from dacitic ejecta from melt inclusion and phase relation studies (Wallace & Anderson, 2000). The estimated water pressure of 1·9 ± 0·3 kbar satisfies the requirement for amphibole stability; requiring a PH2O of more than 1·3 kbar at c. 850°C (e.g. Rutherford et al., 1985). A water pressure of 1·9 kbar indicates a felsic magma reservoir depth of c. 7 km.
Slow decompression of the lower fall deposit felsic magma
The lower fall deposit felsic magma experienced exceptional heating just prior to eruption (Fig. 17). The outer zones of plagioclase phenocrysts certainly crystallized during and after the heating event, because FeO and MgO in those parts are enriched compared with the inner parts of the crystals (Figs 11b and 12). Given that plagioclase growth in the lower fall felsic magma ended at the felsic reservoir (1·9 ± 0·3 kbar), similar to other felsic magmas, the An content of the plagioclase rims in
TOP
ABSTRACT
INTRODUCTION
