JOURNAL OF PETROLOGY
VOLUME 48
NUMBER 8
PAGES 1543^1567
2007
doi:10.1093/petrology/egm029
Remobilization of Highly Crystalline Felsic
Magma by Injection of Mafic Magma:
Constraints from the Middle Sixth Century
Eruption at HarunaVolcano, Honshu, Japan
YUKI SUZUKI1* AND SETSUYA NAKADA2
1
INSTITUTE OF MINERALOGY, PETROLOGY AND ECONOMIC GEOLOGY, GRADUATE SCHOOL OF SCIENCE, TOHOKU
UNIVERSITY, AOBA-KU, SENDAI, 980-8578, JAPAN
2
EARTHQUAKE RESEARCH INSTITUTE, UNIVERSITY OF TOKYO, 1-1-1, YAYOI, BUNKYO-KU, TOKYO, 113-0032, JAPAN
RECEIVED JULY 21, 2006; ACCEPTED MAY 15, 2007
ADVANCE ACCESS PUBLICATION JUNE 22, 2007
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.
I N T RO D U C T I O N
KEY WORDS: high-crystallinity felsic magma; magma plumbing
system; multistage magma mixing; upward dragging of injected
magma; vent opening by low-viscosity magma
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 & Martı́,
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).
*Corresponding author. Telphone: þ81-22-795-5786. Fax: þ81-22795-7763. E-mail: [email protected]
ß The Author 2007. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oxfordjournals.org
JOURNAL OF PETROLOGY
VOLUME 48
NUMBER 8
AUGUST 2007
Fig. 1. Distribution of ejecta from the Futatsudake eruptions and earlier ejecta of Haruna volcano (after Oshima (1986) and Soda (1989)).
Earlier ejecta include those of all stages (Stage 1^5) in Haruna activity. Inset shows the location of Haruna volcano in the NE Japan arc.
Samples of the middle sixth century eruption were collected at Owazawa and Ohinata. Only fall deposits are observed at Ohinata. Py-flow,
pyroclastic flow; VF (inset), volcanic front.
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 endmember 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
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(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.
B AC KG RO U N D O F F U TAT S U DA K E
VO L C A N I S M
Haruna volcano is located in the southern end of the
northeast Japan arc (Fig. 1). The activity of this volcano
(4300 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 SiO2^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 (501km3)
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; 01km3 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 13 km3 and
03 km3, respectively (Soda,1996).
SAMPLES
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).
Petrography
Crystals in pumice and dome lava include phenocryst
phases and microlites in the groundmass. The phenocryst
phase can be classified as ‘phenocryst’ (4300 mm across)
and ‘microphenocryst’ (100^300 mm 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 520 mm across). Amphibole
exists, but not as a microlite (Table 1).
The common phenocryst phases for all ejecta are plagioclase (52 mm), orthopyroxene (52 mm), amphibole
(55 mm), magnetite (51mm) and ilmenite (5700 mm)
(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 (52 mm) and
microphenocryst in the dome lava (Table 1; Fig. 4), and is
found as a microphenocryst (100 mm across) in the gray
pumice and the gray part of the banded pumice (Table 1;
Fig. 4). Clinopyroxene is found as a phenocryst (5500 mm;
Fig. 3d) and as a microphenocryst in the dome lava, and as
a microphenocryst (5300 mm; 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.
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 mm wide).
Some orthopyroxene grains in the dome lava, gray
pumice and gray part of banded pumice have clinopyroxene reaction rims (20 mm). Some of the grains of amphibole are mantled by aggregations of pyroxene, plagioclase,
Fe^Ti oxide and glass (up to 100 mm 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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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.
Fig. 3. Photomicrographs of representative ejecta from the middle sixth century eruption at Futatsudake (plane-polarized light). All scale bars
represent 1mm. (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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Table 1: Phenocryst and groundmass phases
Phenocryst and microphenocryst
Dome lava
Ol
Cpx
Am
Opx
Pl
Mt
Il
Groundmass
P (Opx)
P
P (aggregate)
P (Cpx*)
P (dusty*)
P
P (exsolution)
Gl þ Pl þ Cpx þ Opx þ Ox
Flow
Top 0–30 cm
—
—
P
P
P
P
P
Gl
Top 30–60 cm
—
—
P (aggregate*)
P
P
P
P
Gl Pl Py Ox
Lower
—
—
P
P
P
P
P
Gl Pl Py
Fall
Upper
—
—
P (aggregate*)
P
P
P
P
G þ Pl þ Py þ Ox
Middle
—
—
P
P
P
P
P
Gl Pl Py Ox
Lower (whitey)
—
P MPh
P (aggregate*)
P
P (dusty*)
P
P
Gl
Lower (grayy)
P MPh(Opx)
—
P (aggregate*)
P (Cpx*)
P (dusty*)
P
P (resorbed)
Gl þ Pl þ Cpx þ Opx þ Ox
Parentheses show reaction rim and disequilibrium texture; P, present; —, absent; MPh, only microphenocryst; Ol, olivine;
Cpx, clinopyroxene; Am, amphibole; Opx, orthopyroxene; Pl, plagioclase; Mt, magnetite; Il, ilmenite; Gl, glass; Py,
pyroxene; Ox, Fe–Ti oxides; aggregate, Pl þ Py þ Ox þ Gl; dusty, dusty zone.
*Not found in all crystals.
y
Includes white and gray parts in banded pumices.
A N A LY T I C A L M E T H O D S
Fig. 4. Backscattered electron image of typical olivine (Ol) in ejecta
from the middle sixth century eruption at Futatsudake. Scale bars represent100 mm. (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.).)
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 mm 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 inTohoku 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 mm, 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
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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
Eruption phase: Lower-Fa
Sample no.:
Middle-Fa Upper-Fa FlT 30–60 FlT 0–30 Lava dome
Gray-1 Gray-4 Gray-5 Banded-3 Banded-6 White-2 White* White-1y
White-3
White-5
White-1 Odake-1 Medake-4 Magodake-2
Major elementsz (wt %)
SiO2
5847
5806
5795
5823
5746
6099
6065
6136
6066
6107
6124
6037
5997
TiO2
058
062
059
063
062
050
054
052
054
052
050
054
054
5979
055
Al2O3
1826
1786
1848
1825
1847
1805
1773
1786
1789
1784
1783
1786
1789
1783
FeO
732
796
740
753
770
622
665
628
672
644
623
655
676
689
MnO
015
017
015
015
016
015
015
015
016
015
015
014
015
015
MgO
363
392
369
359
386
290
310
273
302
284
284
313
325
340
CaO
807
796
825
810
832
744
744
729
733
735
737
767
772
772
Na2O
288
283
288
288
282
304
303
306
295
305
311
304
302
298
K2O
055
053
053
055
051
064
063
067
065
065
065
062
061
059
P2O5
009
009
009
010
010
008
009
009
009
009
008
009
009
009
Trace elements (ppm)
Nb
3
3
3
3
2
4
3
3
2
4
4
4
3
3
Zr
49
48
48
50
49
55
57
56
55
57
56
54
53
54
Y
11
8
8
9
10
16
5
6
13
7
7
11
10
9
Sr
283
274
287
283
286
270
280
272
280
275
260
280
284
280
Rb
16
17
17
17
15
20
17
19
16
18
21
17
18
15
Zn
70
78
73
75
75
67
64
66
70
65
64
66
70
71
Cu
36
32
35
32
35
13
6
2
5
4
5
27
35
26
Ni
9
7
8
8
8
1
2
1
4
2
1
6
6
6
Co
25
23
24
21
24
18
17
17
18
18
17
20
21
20
Cr
11
12
16
13
15
7
3
4
6
5
5
12
17
16
V
179
173
165
182
177
134
147
135
145
140
138
142
167
161
Sc
21
23
25
24
24
18
20
24
20
19
16
17
18
22
Ba
131
130
121
168
133
168
176
154
177
140
179
155
172
205
All data for fall and flow deposits are from the Owazawa site (Fig. 2). (For eruption phases and pumice types (gray,
banded, white), see text.) Fa, fall; FlT, flow top.
*Mixture of several pumice fragments.
y
Used as a representative bulk composition of white pumice in Tables 5 and 7.
z
All oxide values are normalized to 100% and total iron is given as FeO.
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.
A N A LY T I C A L R E S U LT S
Whole-rock chemistry
(575^604 wt %) than the white pumice. The SiO2 content
of the banded pumice (575^604 wt %) is similar to that of
the gray pumice (579^585 wt %) or higher. The SiO2 content of the dome lava (597^604 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^SiO2, K2O^Cr), however, the trend for gray and
banded pumices is different from that of the dome lava.
Mineral chemistry
The SiO2 content of the white pumice is nearly constant,
independent of the eruptive unit (600^616; Fig. 2).
Gray and banded pumices have lower SiO2 content
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
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Fig. 5. Variation of SiO2, TiO2 and CaO vs MgO (wt %), and Zr and Cr (ppm) vs K2O (wt %) of whole-rock data.
Clinopyroxene
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.
Phenocrysts and microphenocrysts in the dome lava
have cores of Mg-number 80^75 and Wo [Ca/
(Mg þ Fe þ Mn þ 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.
Olivine
Orthopyroxene
Olivine crystals all show normal zoning in Mg-number
[100 Mg/(Mg þ 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 Mgnumber vs NiO diagram, the composition of olivine in
dome lava overlaps with that of the olivine microphenocrysts from the lower fall deposit (Fig. 6b).
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).
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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 þ 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.
Al(IV), both of which change in a sub-parallel fashion.
The rims are relatively low in Na þ K (Fig. 10) and
Al(IV). No systematic chemical changes were found
between eruption units.
Plagioclase
Fig. 7. Variation of Mg-number in cores and rims of clinopyroxene
phenocrysts and microphenocrysts in dome lavas.
The core composition range (Mg-number 624^657 and
Wo 08^18) is uniform throughout the eruption. Rims of
Type 1 have Mg-number 653^675 and Mg-number
653^678 for the gray and white parts of pumices
(including banded pumices), respectively, for lower fall,
and Mg-number 660^732 for the dome lava. These
values are higher than those for others (Mg-number
638^660). 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).
Amphibole
Most amphibole phenocrysts are ‘hornblende’ (Fig. 10).
They have oscillatory zoning in Na þ K (in A site) and
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 Ca/(Ca þ Na)] (Figs 11 and 12).
The FeO and MgO contents are relatively homogeneous
with 02^04 wt % and 005 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 05 wt % and c.
40% at 005 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).
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
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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.
Fig. 9. Mg-number vs Wo [Ca/(Mg þ Fe þ Mn þ 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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of the ulvo«spinel 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.
Ilmenite
Cores of ilmenite phenocrysts are chemically homogeneous, with x-ilm 055^060 (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
Fig. 10. Si vs Na þ K (in A site) diagrams for amphiboles in representative samples. Dotted line indicates the boundary between tschermakite (Ts) and hornblende (se¤ nsu strı´cto) (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.
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 (An50 and An70,
respectively).
Magnetite
The cores of magnetite phenocrysts are chemically homogeneous, with x-usp 017^020 (where x-usp is the fraction
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 820^8508C and
log fO2 of 103 to 109 for the core pairs from all samples and the rim pairs without compositional changes from
the cores. The algorithm yields a temperature of 8808C
and log fO2 of 99 to 108 for the rim pairs from the
white pumice fraction (including banded pumice) of the
lower fall deposit.
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 mm 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 SiO2 content of the groundmass for the white
pumice is in the range of 774^784 wt % (Table 3); no
systematic change between eruption units was found.
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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. 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.
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.)
The groundmass of the dome lava is less fractionated
(SiO2 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 (370^488%) and to the gray pumice
(305^354%) (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 06 vol. %
in total for the dome lava (Table 4).
DISCUSSION
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
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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.
Table 3: Representative groundmass compositions
White pumice
Lower fall
Middle fall
Flow top 0–30 cm Odake Magodake
Sample GMS* Gl
GMS* Gly
GMS* Gl
SiO2
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.)
Dome lava
7812 (025)z 7835 (027) 7744 (038)
TiO2
031 (003)
Al2O3
1245 (013)
024 (003)
GMS
GMS
7300
718
023 (005)
033
030
1282 (021) 1288 (024)
1374
1310
FeO*
171 (011)
159 (009)
158 (008)
296
341
MnO
007 (004)
006 (003)
009 (002)
007
010
MgO
045 (001)
038 (003)
042 (002)
143
301
CaO
237 (002)
244 (010)
246 (007)
426
420
Na2O
321 (010)
295 (013)
359 (008)
326
310
K2O
131 (007)
118 (003)
132 (004)
095
098
Original 9967 (064)
9657 (211) 9884 (151)
—
—
—
—
total
n
are identical to the white pumice and gray pumice of the
lower fall deposit. We stress the following properties:
(1) 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);
5
6
5
All oxide values are normalized to 100% and total iron is
given as FeO. n, number of analyses for glass; GMS,
groundmass; Gl, glass.
*Crystal-free groundmass.
y
Used in Tables 5 and 8.
z
Standard deviation. Not shown for dome lava, because
data averaged to yield the composition in this table are
from various GMS phases (see text for details).
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Table 4: Phenocryst phase volume proportions in gray pumice and dome lavas
Occurrence:
Lower fall gray pumice
Dome lava
Sample:
No. 1
No. 4
No. 5
Odake-1
Odake-6
Medake-3
Magodake-4
Bulk composition:
585, 36*
581, 39
579, 37
604, 31
598, 33
600, 33
598, 34
501
501
501
vol. %
Olivine
Clinopyroxene
01
02
02
01
02
01
03
03
Plagioclase
258 (729)
231 (736)
218 (715)
332 (703)
399 (823)
275 (749)
363 (793)
Amphibole
50 (141)
38 (121)
32 (105)
84 (177)
43 (88)
45 (124)
40 (87)
Orthopyroxene
39 (110)
36 (115)
43 (141)
41 (86)
29 (60)
38 (104)
42 (92)
Fe–Ti oxides
07 (20)
09 (29)
12 (39)
16 (34)
14 (28)
08 (23)
13 (29)
Total phenocryst
354
314
305
475
488
370
463
Groundmass
646
687
695
525
512
630
537
Phenocryst phases includes both phenocrysts and microphenocrysts. Volume percent values are based on point counts up
to about 2000. Data in parentheses are volume percent plagioclase, amphibole, orthopyroxene and Fe–Ti oxides, as a
proportion of the total phenocryst population.
*SiO2 and MgO contents (wt %), in order.
(4) enrichment in FeO and MgO at and near the rims of
plagioclase (Figs 11a, b and 12);
(5) 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).
Table 5: Phenocryst content in white pumice
Case A (Pl An66)
Case B (Pl An88)
wt %
vol. %
wt %
vol. %
Pl (2700)
369
363 (719)
353
347 (708)
Am (3200)
114
94 (185)
128
105 (214)
Opx (3500)
37
29 (57)
26
20 (41)
Mt (5200)
31
16 (31)
33
17 (34)
Il (4700)
07
04 (08)
02
01 (02)
Total phenocryst
557
505
543
490
GMS-Gl (2400)
443
495
457
510
Mass-balance (mixing) calculations were used to estimate
proportions of phenocryst phases and groundmass for
white pumice (middle fall). Bulk-rock and groundmass
compositions are listed in Tables 2 and 3, and the core
compositions of phenocrysts used are the averaged values
for all eruptive units except for An66 and An88 of
plagioclase. Density values (kg/m3) in parenthesis in the
first column are from Deer et al. (1992) and Hall (1987).
An, An mol %; Pl, plagioclase; Am, amphibole; Opx,
orthopyroxene; Mt, magnetite; Il, ilmenite; GMS, groundmass; Gl, glass. Volume percents among phenocrysts are
shown in parentheses in the forth and seventh columns.
(2) reversely zoned orthopyroxenes in the white parts of
the pumice (including the banded pumice) (Figs 8
and 9);
(3) chemical disequilibrium between olivine and orthopyroxene in the gray parts of the pumice (including
the banded pumice) and the dome lava (Fig. 16a);
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 fO2 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 fO2
(Hattori & Sato, 1996). However, a simultaneous increase
of MgO does not support a change in fO2. 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
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Fig. 16. Mg^Fe distribution between cores of olivine and pyroxene.
(a) olivine Mg-number vs orthopyroxene Mg-number; (b) olivine
Mg-number vs clinopyroxene Mg-number. Data are plotted for both
phenocrysts and microphenocrysts. Each diagram shows the average
with the compositional range. Ranges of equilibrium were calculated
from KD [(Fe/Mg)crystal/(Fe/Mg)melt] of 027^033 for olivine
(Roeder & Emslie, 1970; Ulmer, 1989), 022^031 for clinopyroxene
(Sisson & Grove, 1993a, 1993b) and 0245^0323 for orthopyroxene
(Beattie, 1993).
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.
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 mm) 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 mm) 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 þ 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).
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 fO2. The loss of H2 gas resulting from devolatilization causes oxidation (e.g. Czamanske
& Wones, 1973), which is likely to occur during syneruptive
degassing processes. However, an increase of f O2 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 (8508C), the diffusion coefficient of Mg in
plagioclase with An70 [ ¼13 1018 m2/s, calculated using
the data of Costa et al. (2003)] is found to be larger than
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the Fe2þ^Mg interdiffusion coefficient in orthopyroxene
(20 1020 m2/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
NUMBER 8
AUGUST 2007
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 mm across)
would be homogenized in 15^12 years at 800^9008C
(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 017^020; Fig. 14), surrounded by
reversely zoned rims that are as thin as 55 mm. 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 15^20 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 Soufrie're
Hills magma became zoned within 2^10 days after temperature changed from 835^850 to 860^8808C. 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
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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.
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
Table 6: Estimated bulk-rock compositions for mafic
magmas
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 inTable 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 031^036 for the two compositions of
aphyric mafic magma (A and B in Table 6). We used
01 for Fe3þ/(Fe2þ þ Fe3þ) 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 03 003 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
(a) Aphyric
(b) Phyric
(c) Basalt in early
Haruna activity
A
B
A
B
Major elements* (wt %)
SiO2
5186
5265
5146
5221
5104
TiO2
074
073
070
070
107
5333
120
Al2O3
1889
1880
1761
1763
1843
1868
FeO
1010
980
1015
987
1064
949
MnO
017
017
017
017
014
016
MgO
575
550
715
680
650
523
CaO
969
949
1016
994
932
926
Na2O
247
252
226
233
243
237
K2O
027
030
025
028
022
025
P2O5
010
010
009
009
021
002
Estimation methods are described in the text.
(a) Averaged chemical composition of aphyric mafic
magma was calculated by extension to 0% of phenocrysts
from the white pumice (Middle-Fa White-1 in Table 2) with
different phenocryst contents, A and B (495 and 510 vol.
%) in Table 5, through three gray pumice samples (Gray-1,
-4 and -5 in Table 2) with phenocryst contents in Table 4.
(b) Chemistry of phyric mafic magma was estimated in the
mass balance shown in Table 7. (c) Ejecta of Stage 1
(Oshima, 1983).
*All oxide values are normalized to 100% and total iron is
given as FeO.
Table 7: Mass-balance calculation to form average dome lava composition
Result of mass balance (wt %)
F.M.
Vol. % of phases in phyric mafic magma
Aphyric M.M.
Crystal in phyric M.M.
A
B
Ol
Cpx
R2
Melt
Ol
Cpx
Case A
868
121
—
03
08
001
933
18
49
Case B
857
—
132
03
08
001
938
17
45
Results of mass-balance calculations; mixed magma ¼ felsic magma (F.M.) þ phyric magma, where phyric magma ¼
aphyric mafic magma (M.M.) þ olivine (Ol) þ clinopyroxene (Cpx). The average composition of dome lavas and a
representative white pumice sample, Middle-Fa White-1 (Table 2), were used as mixed and felsic magmas, respectively.
Aphyric mafic magmas of A and B in Table 6 were tested (Cases A and B, respectively). Average compositions of olivine
and clinopyroxene were used. Densities used for melt, olivine and clinopyroxene (in kg/m3) were 2650, 3400 and 3300,
respectively (Hall, 1987).
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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 inTable 6. (b) Mafic magmas are plotted, based on estimated
K2O contents (Table 6, average of A and B).
85^86:12^13: 03: 03 (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).
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
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 þ 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.)
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 (Mgnumber of 77^79 and 77^80, respectively) are in the composition range of the dome lavas (Figs 6a and 7).
The calculated liquidus temperature (1130^11508C) shows
no correlation with pressure. The clinopyroxene barometer
(Nimis & Ulmer, 1998) using phenocryts in the dome lavas
yields 3^4 kbar pressure at 11308C. 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).
The estimated densities of melt, olivine and clinopyroxene under the above conditions are 2600 kg/m3,
3400 kg/m3 and 3300 kg/m3. 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/m3(e.g. granite^diorite, Best & Christiansen,
2001), a pressure of 4 kbar would correspond a depth of
c. 15 km (Fig. 20).
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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 53 04 wt % (n ¼14). This corresponds to the
solubility of water in the felsic melt at 19 03 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
19 03 kbar satisfies the requirement for amphibole stability; requiring a PH2 O of more than 13 kbar at c. 8508C
(e.g. Rutherford et al., 1985). A water pressure of 19 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 (19 03 kbar), similar to other felsic
magmas, the An content of the plagioclase rims in the
lower fall felsic magma should be higher compared with
other felsic magmas, because of the heating event.
However, if we focus on plagioclase microphenocryst
rims, which provide a wider domain for last-stage growth,
we find that the plagioclase rims of the lower fall deposit
magmas have similar or lower An contents compared
with those from other eruption phases (Fig. 13). If the rims
are relatively low in An content (e.g. LFA-5 in Fig. 13), that
part (low-An overgrowth) is clearly distinguished from the
inner part. The above observations suggest that the lower
fall felsic magma underwent syneruptive phenocryst crystallization probably as a result of the slow decompression.
The mechanism of increased crystallization in slowly
decompressed magma is due to the time lag between the
physicochemical change and the crystallization response
(e.g. Lasaga, 1981). Application of the plagioclase hygrometer (Housh & Luhr, 1991) to the lowest An content rim
(An50) of the plagioclase phenocrysts in the lower fall
white pumice (Fig. 13; Table 3) yields 35 wt % H2O
(Ab model basis) and 20 wt % H2O (An model basis) at
8808C, at a reservoir pressure of 19 kbar or less. Given that
the reservoir magma was saturated with water, the water
content indicates the crystallization of the An50 rims at
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c. 05 kbar. This represents the final pressure of slow decompression (Fig. 20 step 3), assuming equilibrium. The width
of the low-An overgrowths (5 mm), which are observed on
the rims of the lower fall deposit plagioclase, indicates
that decompression occurred just before eruption; an overgrowth of 5 mm can be formed within 6 days after fast
decompression to 05 kbar in a dacitic melt at 9008C
(Suzuki et al., 2007).
Decompression can be achieved by either magma and/or
volatile discharge from the reservoir (e.g. Druitt & Sparks,
1984) or by magma ascent. In the former case, the reservoir magma may experience general decompression,
resulting in the formation of low-An plagioclase rims
throughout the eruptive materials, which is not found in
the present case. Thus, the slow decompression, favored
here, corresponds to the slow ascent to the surface
(Fig. 20, step 3). Only magma that erupted earliest (i.e.
the lower fall deposit) can ascend slowly, compared with
succeeding magma batches (Fig. 20, step 3), probably
because it takes time to form or initially open the conduit
or vent.
Clinopyroxene microphenocrysts found only in the
white pumices and the white parts of the banded pumices
of the lower fall deposit (Fig. 3b; Table 1) could have been
stabilized upon decompression. The clinopyroxene liquidus
temperature increases in the felsic magma with progressive
decompression (Rutherford et al., 1985).
Eruption mechanisms
Physical properties of the magmas
In this section we focus on the physical properties of the
magma in the felsic reservoir, and specifically where the
felsic magma had an interface with mafic magma. By
taking into account the heating of the felsic magma
(lower fall deposit) by the mafic magma, we estimated
the physical properties of the felsic magma before and
after the heating (Table 8). To calculate melt viscosity, we
used the formulation for a water-bearing melt of Shaw
(1972). The effect of phenocryst presence can be introduced
by multiplying the melt viscosity by (1 ^ a)n (Marsh, 1981),
where a and n are constants [a ¼16, n ¼ 25, is crystallinity (5 1)]. The viscosity of the felsic magma at 8808C is
(10^15) 108 Pa/s (Table 8). The maximum viscosities of
the aphyric and phyric mafic magmas are 72 101 Pa/s
and (95^97) 101 Pa/s, respectively. Thus, the viscosity of
the felsic magma is about seven orders of magnitude larger
than that of the mafic magma.
The density of the mafic magma decreased little as a
result of vesiculation before its injection into the felsic
magma reservoir, because the mafic reservoir magma was
undersaturated with water (2 wt % H2O at 4 kbar) and
vesiculation would never occur unless the magma is
decompressed to 05 kbar (Fig. 19). The densities of the
aphyric and phyric mafic magmas are 2530^2540 kg/m3
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Table 8: Physical conditions of end-member magma
Felsic magma
Temperature (8C)
Mafic magma
No-heating
Heating*
820–850y
880y
SiO2 wt % in meltz
784
H2O wt % in meltx
53
Phenocryst (vol. %)ô
Aphyric
Phyric
1130y–1150
519–527y
20
490–505
0
62–67
Viscosity (Pa/s)
(20–29) 108
(10–15) 108
72 101
(95–97) 101
Density (kg/m3)
2520
2510–2520
2530–2540
2580–2590
Viscosities for phyric magmas are variable depending on abundances of phenocrysts. Partial molar volumes of oxides
and H2O (Burnham & Davis, 1971; Lange, 1994) were used for calculation of melt density. Densities in Table 5 and
3400 kg/m3 for olivine and 3300 kg/m3 for clinopyroxene were assumed based on Fig. 19. (See the text for details.)
*Phenocryst contents and melt compositions used are the same as those in the no-heating case, because decompression
after heating masked the values at heating.
yUsed for minimum and maximum viscosities for felsic and mafic magmas, respectively.
zRefer to Tables 3 and 6a.
xValues in reservoir condition (19 kbar).
ôRefer to Tables 5 and 7.
and 2580^2590 kg/m3, respectively (Table 8). Moreover, the
density of the felsic magma is similar to that of the mafic
magmas, because of its high crystallinity (about 50 vol.
%).
Necessary conditions for mixing
In the present case, mixed magmas (gray pumice, gray
part of banded pumice and dome lava) appear to have
been formed through complete hybridization (Sparks &
Marshall, 1986). Sparks & Marshall (1986) showed that
hybridization usually requires a small temperature difference between the end-member magmas and a high proportion of the high-temperature magma (450 wt %). In the
scenario we envisage at Futatsudake, the mixing ratio of
the mafic magma is 29^40 wt % for the gray pumices, as
discussed above, and 13^14 wt % for the dome lavas
(Table 7), clearly smaller than this value.
We propose a multistage mixing model for the formation
of the mixed magmas at Futatsudake, considering the
coexistence of phenocrysts with different zoning patterns
such as orthopyroxene in Fig. 8. The product of mixing is,
subsequently, mixed further with the felsic magma.
First, mafic magma in a higher proportion than that estimated from the final product was mixed with the felsic
magma, producing a homogeneous magma with
intermediate temperature between the mafic and felsic
magmas. This mixing product is mixed again with the
remaining felsic magma. In this scenario, phenocrysts
with different histories (zoning profiles) can coexist in the
final products. First, phenocrysts that originated from the
mafic magma show normal zoning. Second, phenocrysts
from the felsic magma that experienced two-stage mixing
show strong reverse zoning in their inner parts (Type 2 opx
in Fig. 8), whereas those experienced one-stage mixing are
zoned slightly (Type 1). It is accepted that second-stage
mixing involves mingling (mafic inclusion formation), if
the disaggregation of mafic inclusions (Feeley & Dungan,
1996; Clynne, 1999) had proceeded completely. No trace of
mafic inclusions has been found in the ejecta of the present
study.
Assuming that the temperature of the mixed magma
was broadly proportional to the mixing ratio of the mafic
magma, the ratio for the initial mixing is found to be
higher in the gray pumice than in the dome lava. Inner
rims of Type 2 orthopyroxene in the gray part of the
pumice and dome lava record temperatures just after the
first mixing event. These orthopyroxene domains are
inferred to have crystallized simultaneously with clinopyroxene that is overgrown on Type 2 orthopyroxene in the
gray part of the pumice, and the rims of clinopyroxene
phenocrysts in the dome lava. This is based on the observation that the compositions of these opx^cpx pairs fall on
isothermal lines of similar temperature in the pyroxene
quadrilateral (Lindsley, 1983), in each gray part of the
pumice and dome lava (Fig. 21). The Wells (1977) twopyroxene geothermometer yielded temperatures of
1100^11508C and 1000^10508C for clinopyroxene and
orthopyroxene pairs from the gray part of the pumices
and the dome lava, respectively.
Possible pre-heating of the felsic reservoir magma
As discussed above, ejecta from the middle sixth century
eruption records two types of heating. The lower fall
white pumices record heating associated with mafic
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Fig. 21. Pyroxene phenocryst rim compositions in the Di^En^Fs^Hd
quadrilateral, to confirm their equilibrium just after the first mixing.
Isothermal lines are based on Lindsley (1983). Clinopyroxene compositions of the dome lava are for the phenocryst rims, whereas those of
the lower fall gray pumice and gray part of banded pumice are for
overgrowth rims on orthopyroxene phenocrysts. Orthopyroxene compositions are for inner rims of Type 2 phenocrysts.
magma injection into the felsic magma reservoir and resultant mixed magma formation. White pumices from the
middle eruption (middle fall to flow top 30^60 cm) probably record heating at a different stage (pre-heating), as
observed only in reversely zoned orthopyroxene rims.
This pre-heating is not related to mafic magma injection
and mixed magma formation, because the white pumices
during the middle eruption phase did not erupt together
with mixed magma (Fig. 17). The pre-heating might have
helped to remobilize the erupted felsic magma, as heating
decreases the viscosity of the magma.
Eruption sequence
The middle sixth century eruption was triggered by the
input of mafic magma into the felsic magma reservoir
(Fig. 20, step 1). Heterogeneity in phenocryst abundance
(aphyric and phyric parts) in the mafic magma might
have been established before its input (Fig. 20, step 1).
Cooling and/or decompression of the mafic magma
NUMBER 8
AUGUST 2007
progressed, resulting in crystallization of small olivines
found in the gray pumices and the gray parts of the
banded pumices (Fig. 4b). The lower fall ejecta (white,
gray and banded pumices) indicate that each of the gray
and white parts is homogeneous, independent of their
dimensions. We envisage that this scenario may occur if
the heated felsic magma (white part) and mixed magma
(gray part) were formed and then stored separately, followed by their mixing with each other just before the
eruption.
It should be noted that the eruption of the mixed
magma (gray pumice and the gray part of the banded
pumice) preceded that of the felsic magma of the middle
fall to flow top 0^30 cm. Based on experimentation,
Snyder & Tait (1996) proposed that convection of overlying
magma, driven by heating from newly injected hot
magma, can result in partial entrainment of the injected
magma (Fig. 20, step 2). In the case of Futatsudake, the
felsic magma seems too viscous to convect immediately
after the mafic magma injection because the crystallinity
of the felsic magma (50 vol. %; Table 5) is close to the
critical crystallinity (55 vol. %) above which magma rheologically behaves as a solid (Marsh, 1981). Thus, it is considered that the mafic magma first intruded into the felsic
reservoir, forming a dyke (Fig. 20, step 1). Following the
intrusion, the felsic magma surrounding the mafic
magma was reheated, becoming less viscous, and then
was eroded by disaggregation into the mafic magma. The
erosion was enhanced not only by heat transfer but also by
dyke propagation, and accompanied by mixing between
mafic and felsic magmas. Once a considerable volume of
mobile felsic magma had accumulated and begun to convect, the upward entrainment of the mafic magma may
have proceeded as suggested by Snyder & Tait (1996)
(Fig. 20, step 2). The similar densities of the mafic and
felsic magmas studied here (Table 8) could have enhanced
the entrainment process. Upward entrainment could have
been localized at the center of the reservoir (Fig. 20, step 2).
If the entrainment process had continued, mixed
magma should have been erupted not only at the beginning of the eruption but also throughout. However, the
occurrence of mixed magma is limited to the beginning
and end of the eruption. Therefore, we infer that entrainment did not continue in the Futatsudake reservoir (Fig. 20,
step 3). Snyder & Tait (1996) showed that a high injection
rate was needed to maintain convection and entrainment.
Takeuchi & Nakamura (2001) showed that the viscous
crystal-rich felsic magma (107 Pa/s) of Komagatake volcano
did not erupt at first, and the mixed magma (104 Pa/s)
could have been as a consequence of dyke propagation.
Eruption of mixed and heated magmas as in the lower fall
deposit may have caused vesiculation of the magma
remaining in the reservoir^conduit system through decompression (Fig. 20, step 4) in the event of a conduit opening
1564
SUZUKI & NAKADA
HARUNA ERUPTION, JAPAN
to the surface, which allowed eruption of the immobile
felsic magma (middle fall^flow top 0^30 cm).
The dome lavas (Fig. 20, step 5) could have been formed
from the phyric mafic magma that remained at the base of
the felsic magma reservoir (Fig. 20, step 3). The eruption of
the mixed (lava dome) magma in the final stages of the
eruption can be explained by any of the following models:
(1) simple turbulent mixing in the conduit during eruption
of the remaining felsic magma and phyric mafic magma;
(2) new injection of mafic magma and its entrainment, as
in the first injection; (3) gravitational overturn of a stratified magma reservoir (through cooling-induced vesiculation of the lower-level phyric mafic magma), leading to
mixing throughout the reservoir. As the transition from
the felsic magma (pumice flow) to the mixed magma
(dome lava) is not continuous, as manifested by the different events, different processes from previous eruption
phases might have occurred in the reservoir just before
the dome lava emplacement. However, the dome lava samples were collected only from the surface of the large dome
complex. Therefore, we cannot preclude the possible
existence of rocks as felsic as white pumice in the lower
parts or interior of the lava dome (Fig. 20, step 5).
CONC LUSION
1. We have investigated the latest Futatsudake eruption,
which occurred in the middle sixth century in the sequence
Plinian fall, pyroclastic flows and lava dome emplacement.
We have revealed the magmatic processes leading to the
eruption of the highly crystalline felsic magma, as well as
the nature of the magma plumbing system beneath the
Haruna volcanic complex (Fig. 20).
2. Gray pumices and the gray parts of banded pumices
from the first Plinian phase (lower fall deposit) and dome
lavas are the product of magma mixing. The mafic endmember switched from aphyric (lower fall) to phyric
(lava dome). The felsic end-member corresponds to the
white pumices found in all eruptive units of the Plinian
and pyroclastic flow phases. The white pumices and white
parts of banded pumices in the lower fall deposits record
heating caused by the adjacent mixed magma and its
mafic end-member.
3. The aphyric mafic magma chemically corresponds to
the melt part of the phyric mafic magma and the two are
closely related genetically. The two mafic magmas are
derived from a storage system at around 15 km depth.
These magmas were formed from an original phyric
magma, through density-segregation of mafic phenocrysts
in the mafic reservoir.
4. The aphyric part of the mafic magma was injected
first into the felsic magma reservoir (c. 7 km depth) without affecting the main part that erupted after the lower
fall deposit and before the lava dome emplacement. This
is because thermal convection in the reservoir, driven by
new magma injection, dragged the injected mafic magma
upward. However, the phyric mafic magma was not
dragged upward and remained at the base of the felsic
reservoir, thereby resulting in the formation of the dome
lavas.
5. Mafic magma injection made the high-viscosity felsic
magma erupt, through the formation of less viscous mixed
and heated magmas (lower fall ejecta) and by the opening
of the conduit and vent by the low-viscosity magmas.
Indeed, the lower fall white pumices have records of syneruptive slow ascent to a depth of 2 km depth, related to
conduit formation.
6. Formation of homogeneous mixed magmas (gray
pumices, gray parts of banded pumices, and dome lavas)
with low mafic contributions requires multistage mixing
(including disaggregation of mafic inclusions), which is
supported by phenocryst zoning patterns.
AC K N O W L E D G E M E N T S
This research was mainly undertaken in the Earthquake
Research Institute, University of Tokyo (ERI) as a part of
the doctoral thesis of Y.S. We are deeply grateful to
Associate Professor A. Yasuda (ERI) for help with the
microprobe work. We would also like to thank Professor
T. Fujii (ERI), Associate Professor T. Ishii, Professor
H. Nagahara and Associate Professor H. Iwamori
(University of Tokyo) for reading the first version in the
thesis. Dr O. Oshima (now retired from University of
Tokyo) is thanked for information on Haruna volcano.
Also, Y.S. thanks Professor T. Yoshida and Associate
Professor M. Nakamura (Tohoku University) for advice.
Dr S. Okumura (Tohoku University) is thanked for lending Y.S. hydrous glass standards. Finally, the manuscript
was greatly improved by the insightful comments from
Dr Dougal Jerram, Dr Charles Bacon, an anonymous
reviewer and Professor John Gamble. Professor John
Gamble also corrected aspects of English language. This
work was partly supported by Grant-in-Aid from MEXT
to S.N. (no. 12304033). Additionally, Y.S. was supported by
The 21st Century COE Program in Tohoku University.
S U P P L E M E N TA RY DATA
Supplementary data for this paper are available at Journal
of Petrology online.
R EF ER ENC ES
Andersen, D. J. & Lindsley, D. H. (1988). Internally consistent solution
models for Fe^Mg^Mn^Ti oxides. American Mineralogist 73, 714^726.
Bacon, C. R. & Hirshmann, M. M. (1988). Mg/Mn partitioning as a
test for equilibrium between coexisting Fe^Ti oxides. American
Mineralogist 73, 57^61.
Beattie, P. (1993). Olivine^melt and orthopyroxene^melt equilibria.
Contributions to Mineralogy and Petrology 115, 103^111.
1565
JOURNAL OF PETROLOGY
VOLUME 48
Berndt, J., Koepke, J. & Holtz, F. (2005). An experimental investigation of the influence of water and oxygen fugacity on differentiation
of MORB at 200 MPa. Journal of Petrology 46, 135^167.
Best, M. G. & Christiansen, E. H. (2001). Igneous Petrology. Oxford:
Blackwell Science.
Blake, S. (1984). Volatile oversaturation during the evolution of silicic
magma chambers as an eruption trigger. Journal of Geophysical
Research 89, 8237^8244.
Blundy, J. D. & Holland, T. J. (1990). Calcic amphibole equilibria and
a new amphibole^plagioclase geothermometer. Contributions to
Mineralogy and Petrology 104, 208^224.
Burnham, C. W. & Davis, N. F. (1971). The role of H2O in silicate melts
I. P^V^T relations in the system NaAlSi3O8^H2O to 10 kilobars
and 10008C. AmericanJournal of Science 270, 54^79.
Carmichael, I. S. E. & Ghiorso, M. S. (1990). The effect of oxygen
fugacity on the redox state of natural liquids and their crystallizing
phases. In: Nicholls, J. & Russell, J. K. (eds) Modern Methods of
Igneous Petrology: Understanding Magmatic Processes. Mineralogical
Society of America, Reviews in Mineralogy 24, 191^212.
Clynne, M. A. (1999). A complex magma mixing origin for rocks
erupted in 1915, Lassen Peak, California. Journal of Petrology 40,
105^132.
Costa, F., Chakraborty, S. & Dohmen, R. (2003). Diffusion coupling
between trace and major elements and a model for calculation of
magma residence times using plagioclase. Geochimica et
Cosmochimica Acta 67, 2189^2200.
Cottrell, E., Gardner, J. E. & Rutherford, M. J. (1999). Petrologic and
experimental evidence for the movement and heating of the preeruptive Minoan rhyodacite (Santorini, Greece). Contributions to
Mineralogy and Petrology 135, 315^331.
Czamanske, G. K. & Wones, D. R. (1973). Oxidation during magmatic
differentiation, Finnmarka Complex, Oslo area, Norway: Part 2,
The mafic silicates. Journal of Petrology 14, 349^380.
Deer, W. A., Howie, R. A. & Zussman, J. (1992). An Introduction to the
Rock-forming Minerals, 2nd edn. Harlow: Longman.
Devine, J. D., Gardner, J. E., Brack, H. P., Layne, G. D. &
Rutherford, M. J. (1995). Comparison of microanalytical methods
for estimating H2O contents of silicic volcanic glasses. American
Mineralogist 80, 319^328.
Devine, J. D., Rutherford, M. J., Norton, G. E. & Young, S. R. (2003).
Magma storage region processes inferred from geochemistry of
Fe^Ti oxides in andesitic magma, Soufrie're Hills Volcano,
Montserrat, W.I. Journal of Petrology 44, 1375^1400.
Druitt, T. & Sparks, R. S. J. (1984). On the formation of calderas
during ignimbrite eruptions. Nature 310, 679^681.
Feeley, T. C. & Dungan, M. A. (1996). Compositional and dynamic
controls on mafic^silicic magma interactions at continental arc
volcanoes: evidence from Cordon El Guadal, Tatara^San Pablo
Complex, Chile. Journal of Petrology 37, 1547^1577.
Folch, A. & Marti, J. (1998). The generation of overpressure in felsic
magma chambers by replenishment. Earth and Planetary Science
Letters 163, 301^314.
Ghiorso, M. S. & Sack, R. O. (1995). Chemical mass transfer in
magmatic processes IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of
liquid^solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology 119,
197^212.
Gill, J. B. (1981). Orogenic Andesites and PlateTectonics. Berlin: Springer.
Hall, A. (1987). Igneous Petrology, 2nd edn. Harlow: Longman.
Hammer, J. E., Cashman, K. V., Hoblitt, R. P. & Newman, S. (1999).
Degassing and microlite crystallization during pre-climactic events
NUMBER 8
AUGUST 2007
of the 1991 eruption of Mt. Pinatubo, Philippines. Bulletin of
Volcanology 60, 355^380.
Hattori, K. & Sato, H. (1996). Magma evolution recorded in plagioclase zoning in 1991 Pinatubo eruption products. American
Mineralogist 81, 982^994.
Holloway, J. R. & Blank, J. G. (1994). Application of experimental
results to C^O^H species in natural melts. In: Carroll, M. R. &
Holloway, J. R. (eds) Volatiles in Magmas. Mineralogical Society of
America, Reviews in Mineralogy 30, 187^230.
Housh, T. B. & Luhr, J. F. (1991). Plagioclase^melt equilibria in
hydrous systems. American Mineralogist 76, 477^492.
Lange, R. A. (1994). The effect of H2O, CO2 and F on the density and
viscosity of silicate melts. In: Carroll, M. R. & Holloway, J. R.
(eds) Volatiles in Magmas. Mineralogical Society of America, Reviews in
Mineralogy 30, 331^369.
Lasaga, A. C. (1981). Implications of a concentration-dependent
growth rate on the boundary layer crystal^melt model. Earth and
Planetary Science Letters 56, 429^434.
Lindsley, D. H. (1983). Pyroxene thermometry. American Mineralogist
68, 477^493.
Machida, H., Arai, F., Oda, S., Endo, K. & Sugihara, S. (1984).
Tephra with special reference to Japanese archeology. In:
Watanabe, N. (ed.) Preservation and Cultural and Natural Science for
Cultural Properties. Kyoto: Dohosha Shuppan, pp. 865^928
(in Japanese).
Marsh, B. D. (1981). On the crystallinity, probability of occurrence,
and rheology of lava and magma. Contributions to Mineralogy and
Petrology 78, 85^98.
Miyashiro, A. (1974). Volcanic rock series in island arcs and active continental margins. AmericanJournal of Science 274, 321^355.
Morgan, G. B., IV & London, D. (1996). Optimizing the electron
microprobe analysis of hydrous alkali aluminosilicate glasses.
American Mineralogist 81, 1176^1185.
Murphy, M. D., Sparks, R. S. J., Barclay, J., Carroll, M. R. &
Brewer, T. S. (2000). Remobilization of andesite magma by intrusion of mafic magma at Soufrie're Hills Volcano, Montserrat, West
Indies. Journal of Petrology 41, 21^42.
Nakada, S. & Motomura, Y. (1999). Petrology of the 1991^1995 eruption at Unzen: effusion pulsation and groundmass crystallization.
Journal of Volcanology and Geothermal Research 89, 173^196.
Nakamura, M. (1995). Continuous mixing of crystal mush and replenished magma in the ongoing Unzen eruption. Geology 23, 807^810.
Nakamura, M. & Shimakita, S. (1998). Dissolution origin and synentrapment compositional change of melt inclusion in plagioclase.
Earth and Planetary Science Letters 161, 119^133.
Nimis, P. & Ulmer, P. (1998). Clinopyroxene geobarometry of magmatic rocks Part 1: An expanded structural geobarometer for anhydrous and hydrous, basic and ultrabasic systems. Contributions to
Mineralogy and Petrology 133, 122^135.
Oshima, O. (1983). Geology, petrology, and mineralogy of Haruna
volcano. Doctoral thesis, University of Tokyo, 280 pp.
Oshima, O. (1986). Haruna volcano. In: Editorial Committee of
Kanto, Part 3 of Regional Geology of Japan (ed.) Regional Geology
of Japan, Part 3 Kanto Region. Tokyo: Kyoritsu Shuppan, pp. 222^224
(in Japanese).
Pallister, J. S., Hoblitt, R. P., Meeker, G. P., Knight, R. J. & Siems, D.
F. (1996). Magma mixing at Pinatubo: petrographic and chemical
evidence from the 1991 deposits. In: Newhall, C. &
Punonhbayan, R. (eds) Fire and Mud: Eruptions and Lahars of Mt.
Pinatubo. Seattle: University of Washington Press, pp. 687^731.
Roeder, P. L. & Emslie, R. F. (1970). Olivine^liquid equilibrium.
Contributions to Mineralogy and Petrology 29, 275^289.
1566
SUZUKI & NAKADA
HARUNA ERUPTION, JAPAN
Rutherford, M. J. & Devine, J. D. (2003). Magmatic conditions and
magma ascent as indicated by hornblende phase equilibria and
reactions in the 1995^2002 Soufrie're Hills magma. Journal of
Petrology 44, 1433^1453.
Rutherford, M. J., Sigurdsson, H. & Carey, S. (1985). The May 1, 1980
eruption of Mount St. Helens: 1, Melt composition and experimental phase equilibria. Journal of Geophysical Research 90, 2929^2947.
Sack, R. O., Carmichael, I. S. E., Rivers, M. & Ghiorso, M. S. (1980).
Ferric^ferrous equilibria in natural silicate liquids at 1bar.
Contributions to Mineralogy and Petrology 75, 369^376.
Sakaguchi, H. (1986). Haji and Sue-ki detected below FA and FP
erupted from Haruna-Futatsudake. In: Educational Board of
Gunma Prefecture (ed.) Reports on the Arato-Kitahara Site, the
Imai-Jinja Kofun Group and the Arato-Aoyagi Site. Maebashi: Gunma
prefecture, pp. 103^119.
Sakaguchi, H. (1993). Methods to infer the age and the season of
volcanic disasters. In: Arai, F. (ed.) Tephro-archeology. Tokyo: Kokon
Shoin, pp. 54^82 (in Japanese).
Scaillet, B. & Evans, B. W. (1999). The 15 June 1991 eruption of Mount
Pinatubo. I. Phase equilibria and pre-eruption P^T^fO2^fH2O
conditions of the dacite magma. Journal of Petrology 40, 381^411.
Shaw, H. R. (1972). Viscosities of magmatic silicate liquids: an
empirical method of prediction. American Journal of Science 272,
870^893.
Sisson, T. W. & Grove, T. L. (1993a). Experimental investigations of
the role of H2O in calc-alkaline differentiation and subduction
zone magmatism. Contributions to Mineralogy and Petrology 113,
143^166.
Sisson, T. W. & Grove, T. L. (1993b). Temperatures and H2O contents
of low-MgO high-alumina basalts. Contributions to Mineralogy and
Petrology 113, 167^184.
Snyder, D. & Tait, S. (1996). Magma mixing by convective entrainment. Nature 379, 529^531.
Soda, T. (1989). Two 6th century eruptions of Haruna volcano, central
Japan. The Quaternary Research (JAPAN) 27, 297^312 (in Japanese
with English abstract).
Soda, T. (1993). Haruna-Futatsudake eruptions in the Kofun period.
In: Arai, F. (ed.)
Tephro-archeology. Tokyo: Kokon Shoin,
pp. 128^150 (in Japanese).
Soda, T. (1996). Explosive activities of Haruna volcano and their
impacts on human life in the sixth century A.D. Geographical
Reports of Tokyo Metropolitan University 31, 37^52.
Sparks, R. S. J. & Marshall, L. A. (1986). Thermal and mechanical
constraints on mixing between mafic and silicic magmas. Journal
of Volcanology and Geothermal Research 29, 99^124.
Sparks, R. S. J., Sigurdsson, H. & Wilson, L. (1977). Magma mixing:
a mechanism for triggering acid explosive eruptions. Nature 267,
315^318.
Suzuki, Y., Gardner, J. E. & Larsen, J. F. (2007). Experimental constraints on syneruptive magma ascent related to the phreatomagmatic phase of the 2000 A.D. eruption of Usu volcano, Japan.
Bulletin of Volcanology 69, 423^444.
Takeuchi, S. & Nakamura, M. (2001). Role of precursory less-viscous
mixed magma in the eruption of phenocryst-rich magma: evidence
from the Hokkaido-Komagatake 1929 eruption. Bulletin of
Volcanology 63, 365^376.
Tomiya, A. & Takahashi, E. (2005). Evolution of the magma chamber
beneath Usu volcano since 1663: a natural laboratory for observing
changing phenocryst compositions and textures. Journal of Petrology
46, 2395^2426.
Ulmer, P. (1989). The dependence of the Fe2þ^Mg cation-partitioning
between olivine and basaltic liquid on pressure, temperature and
composition: an experimental study to 30 kbars. Contributions to
Mineralogy and Petrology 101, 261^273.
Venezky, D. Y. & Rutherford, M. J. (1997). Preeruption conditions and
timing of dacite^andesite magma mixing in the 22 ka eruption at
Mount Rainier. Journal of Geophysical Research 102, 20069^20086.
Venezky, D. Y. & Rutherford, M. J. (1999). Petrology and Fe^Ti oxide
reequilibration of the 1991 Mount Unzen mixed magma. Journal of
Volcanology and Geothermal Research 89, 213^230.
Wallace, P. & Anderson, A. T., Jr (2000). Volatiles in magmas. In:
Sigurdsson, H. (ed.) Encyclopedia of Volcanoes. New York: Academic
Press, pp. 149^170.
Watanabe, Y. & Takahashi, M. (1995). Temporal variations of major
elemental chemistry of the Haruna volcano. In: Abstracts, 1995 Fall
Meeting of the Volcanological Society of Japan, p. 62 (in Japanese).
Wells, P. R. A. (1977). Pyroxene thermometry in simple and complex
systems. Contributions to Mineralogy and Petrology 62, 129^139.
1567