JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
PAGES 2303–2331
2001
Geochemical Evolution of Akagi Volcano,
NE Japan: Implications for Interaction
Between Island-arc Magma and Lower
Crust, and Generation of Isotopically
Various Magmas
KATSURA KOBAYASHI∗ AND EIZO NAKAMURA
PHEASANT MEMORIAL LABORATORY OF GEOCHEMISTRY AND COSMOCHEMISTRY, INSTITUTE FOR STUDY OF THE
EARTH’S INTERIOR, OKAYAMA UNIVERSITY AT MISASA, TOTTORI, 682-0193, JAPAN
RECEIVED APRIL 30, 2000; REVISED TYPESCRIPT ACCEPTED JUNE 10, 2001
Major and trace element, and Sr, Nd and Pb isotopic compositions
were determined for whole-rock samples from the ‘isotopically
anomalous’ Akagi volcano in the volcanic front of the NE Japan
arc. Sr and Nd isotopic compositions of phenocrysts were also
analyzed together with their major and trace element compositions.
Compared with the other volcanoes from the volcanic front, the
whole-rock isotope compositions of Akagi show highly enriched
characteristics; 87Sr/86Sr = 0·7060–0·7088, Nd = −0·40
to −8·6, and 208Pb/204Pb = 38·4–38·8. The rare earth
element (REE) patterns are characterized by heavy REE (HREE)
depletions with U-shaped patterns from middle REE (MREE) to
HREE, suggesting that amphibole fractionation was induced by a
reaction between clinopyroxene and H2O-rich magma in the lower
crust. The integrated isotope and trace elements systematics, and
tectonic structure beneath Akagi volcano, suggest that lower-crustal
assimilation by the H2O-rich primary magma could have been
affected by the double subduction of Philippine Sea and Pacific
oceanic plates. This double subduction could have supplied larger
amounts of water to the magma source region in the wedge mantle
than in the case of a single subduction zone. Significant differences
in isotopic compositions are observed between phenocrysts and the
coexisting melts. Such isotopic disequilibrium may have resulted
from magma mixing between an isotopically depleted aphyric and
an enriched porphyritic magma in a shallow magma chamber.
The geochemical characteristics of these end-member magmas were
retained in the lower crust, despite differing extents of lower-crustal
assimilation by the H2O-rich magmas.
Island-arc volcanism has been considered to result from
partial melting of the wedge mantle induced by addition
of slab-derived fluid-rich materials in subduction zones.
This general model is supported by many geochemical
studies employing major and trace element compositions,
and radiogenic isotope systematics of arc volcanic ejecta
(e.g. Nakamura et al., 1985; Woodhead & Fraser, 1985;
Tatsumi et al., 1986; Ishikawa & Nakamura, 1994; Miller
et al., 1994; Ryan et al., 1995; Shibata & Nakamura,
1997). These previous works mainly focused on the source
characteristics in relation to the evolution of the mantle
wedge in subduction zones. The magmatic evolution of
individual volcanoes, after partial melting of the wedge
mantle, and in shallow magma chambers has received
less attention. Isotope and trace element geochemistry
has typically been restricted to basaltic volcanic rocks,
not the felsic rocks, to avoid the complexities of processes
such as crystal fractionation, magma mixing and crustal
assimilation. It is, however, essential to characterize the
shallower magmatic processes along with the source
characteristics. Magmas erupted at the Earth’s surface
∗Corresponding author.
E-mail: [email protected]
 Oxford University Press 2001
Akagi volcano; H2O-rich magma; isotopic disequilibrium;
lower-crustal assimilation; magma mixing
KEY WORDS:
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 42
possess original source-related characteristics considerably modified by post-melting physico-chemical processes that may obscure the source signature.
In this study, Akagi volcano in the NE Japan arc was
investigated because of its distinctive isotopic composition
relative to neighboring volcanoes from the volcanic front
of the NE Japan arc. This arc is one of the ideal localities
for discussing island-arc magmatism, because the structure of the subduction zone is well defined by abundant
seismological data (e.g. Utsu, 1974; Yoshii, 1979; Zhao
et al., 1990; Zhao, 1992). Moreover, systematic acrossarc variations of chemical and isotopic compositions were
correlated with the depth of the Wadati–Benioff Zone
(Notsu, 1983; Nakamura et al., 1985; Sakuyama & Nesbitt,
1986; Shibata & Nakamura, 1997). These across-arc
systematics have been explained by continuous dehydration of the subducting oceanic slab resulting in a
continuous decrease in the slab component in the source
region with increasing depth to the subducted slab (Shibata & Nakamura, 1997). Volcanoes in the northern part
of the volcanic front show a relatively narrow range of
87
Sr/86Sr ratio (0·7038–0·7045) (Notsu, 1983; Shibata &
Nakamura, 1997). However, the Sr isotope composition
in the southern part of the volcanic front increases to
the south and has the highest 87Sr/86Sr ratio of >0·7087
at Akagi volcano (Notsu, 1983; Kersting et al., 1996).
According to Gust et al. (1997), such systematic isotopic
change along the volcanic front in the NE Japan arc
could be explained by the contribution of subcontinental
lithospheric mantle with some crustal contamination
based on Sr–Nd–Pb isotope systematics and trace element
geochemistry. They also suggested that the primary
magmas of the volcanoes located south of the Tanakura
Tectonic Line might have been involved with IndianOcean-type mantle, although they pointed out the possibility of lower-crustal contamination in the petrogenesis
of the felsic volcanic rocks. However, their data are
limited for individual volcanoes (only one to four samples
without petrological description), making it difficult to
discriminate samples that might have been involved with
crustal contamination. It is, therefore, risky to discuss
their source materials in the mantle beneath individual
volcanic centers based on previous studies with insufficient petrological description of samples or a systematic dataset for each volcano.
From such a point of view, we investigated Akagi
volcano, measuring major and trace element compositions and multi-isotope compositions including Sr,
Nd and Pb isotopes for whole-rock samples, to comprehensively understand the magmatic processes involved
in the formation of the isotopically most anomalous
volcano in the volcanic front of the NE Japan arc, and
to assess the mechanisms of along-arc variation with
respect to tectonic setting. In addition, the trace element,
NUMBER 12
DECEMBER 2001
and Sr and Nd isotope compositions of phenocryst minerals were determined to further understand the shallow
magma chamber processes and thereby elucidate the
genesis of andesite magma in subduction zones.
Geological and petrological features of
Akagi volcano
Akagi volcano (36°33′N, 139°12′E) is located in the
southern part of the volcanic front of the NE Japan arc
(Fig. 1). This area is characterized by the presence of
many Quaternary volcanoes (e.g. Haruna, Hotaka,
Nikko-Shirane volcano), and is one of the most volcanically active regions in Japan. As illustrated in Fig. 1,
an unusual structure exists in the mantle wedge beneath
the volcanic area including Akagi volcano; that is, a
double subduction zone, in which the Philippine Sea
plate is subducted into the wedge mantle above the
Pacific plate (Ishida, 1991). The tip of the Philippine Sea
plate is subducted to depth of >90 km just under Akagi
volcano, and it probably adheres closely to the subducting
Pacific plate, of which the interface depth is >110 km.
Moreover, the depth of the Moho under Akagi volcano
is >39 km, some 4 km deeper than is typical beneath
the volcanic front of the NE Japan arc (Zhao et al., 1990;
Zhao, 1992). These unusual tectonic conditions probably
contribute to the formation of the distinctive geochemical
characteristics of Akagi volcano.
Although Akagi volcano is known to be mostly Quaternary in age, the lack of systematic radiometric ages
precludes a clear understanding of the commencement
of volcanic activity. On the basis of stratigraphical relationships between Akagi volcano (Moriya, 1968) and
the neighboring Komochi volcano (Iizuka, 1996), the
activity of Akagi volcano is thought to have started before
350 ka, and lasted intermittently to 30 ka. The latter age
is based on the only available radiometric age, determined
by the 14C method by Koga (1981).
The volcanic history of Akagi volcano is divided into
three main stages based on the eruption style: the older
stratovolcano formation stage (O) characterized by the
eruption of andesitic lava flow and ejecta of scoria, the
younger stratovolcano formation stage (Y) characterized
by pyroclastics without lava flows, and the central cone
formation stage (Cc) (Moriya, 1968). The older stratovolcano formation stage (O) is further divided into
three substages based on the volcanic stratigraphy: the
early (Oe), middle (Om) and late (Ol) substages. The
main volcanic stages and substages are geologically distinguishable. However, the eruptive sequence of volcanic
ejecta within each stage is unclear based on stratigraphy.
The present volcanic flank is mainly formed of pyroclastics of the Y stage and of secondary deposits. Lava
flows can be observed in the caldera, which formed after
the activity of the Y stage, and near the caldera wall.
2304
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Fig. 1. Geological and sample locality maps for Akagi volcano. The geological map is modified from Moriya (1968). In the inset diagram the
volcanic front of the NE Japan arc and depth contours of the Wadati–Benioff Zone of the Philippine Sea plate (Ishida, 1992) are shown.
Sample locality and description
To investigate the magmatic processes of Akagi volcano,
volcanic ejecta were collected based on the volcano
stratigraphy established by Moriya (1968). The sampling
localities are presented in Fig. 1. Samples were picked
from fresh parts of outcrops to minimize alteration effects,
and examined under the petrographic microscope to
confirm the absence of significant secondary alteration.
Petrographic observations indicate that the Akagi rock
samples are highly porphyritic, with phenocrysts mainly
consisting of plagioclase, hypersthene, augite and opaque
minerals (see Table 1). Some of the felsic samples with
SiO2 contents exceeding 58 wt % in the Ol substage and
Y and Cc stages contain phenocrysts of hornblende
instead of augite, even without quartz phenocrysts. Olivine phenocrysts are absent in the studied samples except
for a sample AK1303, which contains an olivine pseudomorph surrounded by orthopyroxene. The mineral assemblage in the groundmass of the Oe and Om substages
is plagioclase + pigeonite ± augite ± hypersthene. In
contrast, samples from the Ol substage and the Y and
Cc stage samples do not contain groundmass pigeonite.
Sample preparation
In the preparation of whole-rock samples for major, trace
element and isotope analyses, chunks free from surface
alteration were picked. These unaltered chunks were
washed with Milli-Q water using an ultrasonic bath, and
dried. Then, they were ground to fine powders with
grain sizes under 200 mesh using a silicon nitride mortar.
To separate phenocrysts, rock powders obtained using
a disk mill were sieved to collect materials with grain
size ranging from 100 to 200 mesh, as the average grain
size of phenocrysts of plagioclase, orthopyroxene, and
clinopyroxene or hornblende is >500 m. The sieved
fraction was then processed to purify the minerals using
conventional heavy liquid methods followed by an isodynamic separator. The mineral separation was finally
2305
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
DECEMBER 2001
Table 1: Phenocryst assemblages and modal compositions of the studied samples
Sample no.
Pl
Opx
Cpx
Ol
Hb
Mt
15·1
0·9
—
—
1·8
0·5
AK2605
32·5
6·7
3·4
—
tr.
1·7
AK2602
28·4
4·7
1·6
—
1·1
1·2
AK2603W
28·0
3·8
1·8
—
—
1·0
AK2603G
29·5
3·2
1·1
—
—
1·0
AK2603B
33·4
4·6
1·0
—
—
1·0
AK2601
14·1
3·2
0·5
—
—
0·5
AK1202
32·6
5·0
3·4
—
tr.
2·0
AK1108
22·6
2·4
—
—
2·1
0·9
AK1102
31·5
4·1
0·6
—
—
1·2
AK1201
31·2
7·6
0·7
—
—
1·3
AK1012
36·4
9·6
1·9
—
—
1·8
AK-A
35·7
6·8
2·7
—
—
1·6
AK1010
21·8
3·6
0·5
—
—
0·6
AK0910
29·9
8·6
0·7
—
—
1·0
AK1103
30·1
1·3
tr.
—
—
0·3
AK1011
30·1
2·6
tr.
—
—
0·9
AK1004
28·8
0·5
0·1
—
—
0·3
AK1002
32·3
2·6
0·2
—
—
0·2
AK2604
23·0
3·6
2·0
—
—
1·1
AK1302
28·4
1·1
0·4
—
—
0·6
AK1303
19·5
5·7
1·5
tr.
—
1·9
Cc
AK0807
Y
Ol
Om
Oe
Pl, plagioclase; Opx, orthopyroxene; Cpx, clinopyroxene; Ol, olivine; Hb, hornblende; Mt, magnetite; tr., trace (<0·1 vol. %).
accomplished by hand-picking, and the purity is regarded
as being better than 99%.
Analytical methods
All geochemical analyses were carried out at the Pheasant
Memorial Laboratory (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa. Major
element compositions of whole rocks were determined
by X-ray fluorescence (XRF) using a Philips PW-2400
system; details of the analytical procedure are described
elsewhere (Takei et al., in preparation). Trace element
analyses of the whole rocks were performed by inductively
coupled plasma mass spectrometry (ICP-MS) using a
Yokogawa PMS 2000 instrument with a flow-injection
method developed by Makishima & Nakamura (1997).
The analytical reproducibility for trace element analyses
of andesitic samples was <5%, and typically >3% (Makishima & Nakamura, 1997).
The analytical procedures, including the chemical separations and mass spectrometry utilized in this study, are
from Yoshikawa & Nakamura (1993), Makishima &
Nakamura (1991a, 1991b) and Koide & Nakamura (1990)
for Sr, Nd and Pb isotope analysis, respectively. Mass
spectrometry was carried out on Finnigan MAT 261 and
MAT 262 instruments equipped with five Faraday cups,
applying a static-multi-collection mode. Normalizing factors used to correct isotopic fractionation of Sr and Nd
are 86Sr/88Sr = 0·1194 and 146Nd/144Nd = 0·7219,
respectively. The Pb isotope analyses were carried out
with a nearly fixed filament temperature of >1200°C.
Measured ratios of reference materials were 87Sr/86Sr =
0·710233 ± 8 (2m) for 100 ng of NIST SRM 987,
143
Nd/144Nd = 0·511818 ± 13 (2m) for 30 ng of La
Jolla, and 206Pb/204Pb = 16·940 ± 15 (2m), 207Pb/
2306
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Pb = 15·494 ± 16 (2m) and 208Pb/204Pb = 36·709
± 38 (2m) for 100 ng of NIST SRM 981.
Major element compositions of phenocrysts were determined by electron probe microanalysis (EMPA), using
a JEOL JXA-8800R instrument, at Misasa following the
techniques of Iizuka (1996). Trace element analyses of
clinopyroxene, orthopyroxene and plagioclase phenocrysts were carried out employing a Cameca ims 5f
ion microprobe at PML, Misasa, following procedures
described by Nakamura & Kushiro (1997, 1998). Standard materials used for the calibration of trace element
compositions were clinopyroxene from mantle xenoliths
for pyroxene and hornblende analyses, and labradorite
megacrysts for plagioclase analyses. These standards have
been well characterized in terms of homogeneity and
concentrations by both ion microprobe and ICP-MS.
Clinopyroxene phenocrysts in the thin sections were
sputtered with an O− primary beam of >10 nA intensity,
resulting in >10 m beam diameter, and orthopyroxene
and plagioclase with 15–20 nA intensity, resulting in
>15 m beam diameter. Positive secondary ions were
collected by ion counting using an energy offset of −60
V for Si, Rb, Sr, Y and Zr, and of −45 V for other
elements from 4500 V acceleration with an energy bandpass of ±10 V. These operational conditions resulted in
(1−2) × 105 c.p.s. for 30Si secondary ion in the analysis
of phenocrysts. Analytical reproducibilities (RSD 1%,
n = 10) for trace elements in the clinopyroxene standard
were typically >5%, except for Ba (60%), and for the
labradorite standard also typically >5%, except for Zr,
Nb, Sm and heavy rare earth elements (HREE) (20–30%).
204
RESULTS AND DISCUSSION
Major element compositions of whole-rock
samples
Major and trace element compositions of the Akagi
samples are given in Table 2. Selected major element
oxides are plotted against SiO2 content in Fig. 2, classifying the samples into the five stages defined by Moriya
(1968) with reference to the Nasu volcanic zone, which
forms the volcanic front of NE Japan (Kawano et al.,
1961). As shown in Fig. 2, the range of SiO2 contents of
the Akagi samples is between 53·4 and 71·4 wt %,
consistent with those of previous studies, indicating that
basaltic rocks have not been discovered at Akagi (Koga,
1984; Yamaguchi, 1990). TiO2, Fe2O3, MgO and CaO
show negative trends against SiO2, and these trends are
essentially consistent with those of the Nasu volcanic
zone. Two different trends, the relatively steeper trend
formed by the samples in the Oe and Om stages and
that in the Ol, Y and Cc stages, can be recognized in
the Fe2O3 and MgO diagrams, as well as in the Na2O
diagram.
The Al2O3–SiO2 diagram shows a more complicated
feature, with a positive trend for the earlier stages of Oe
and Om, and a negative trend for the later stages of Ol,
Y and Cc. The former positive trend may be explained
by considerable plagioclase accumulation in a shallow
magma chamber, as the deduced Al2O3 contents in the
melts obtained by subtracting the phenocryst compositions from those of the whole rocks form a negative
trend against SiO2 overlapping the samples from the
later stages as shown in Fig. 3. The major element
compositions, after subtraction of the phenocrysts in
Adatara volcano (the neighboring volcano located in the
volcanic front of NE Japan), are also shown based on
the data given by Fujinawa (1988, 1990). The corrected
Al2O3 concentrations for Akagi volcano indicate a similar
negative trend to that of Adatara volcano; however, that
of Akagi is still systematically higher than for Adatara at
similar SiO2 contents. The relatively high content of
Al2O3 in the Akagi samples is consistent with the results of
Yamaguchi (1990). The differentiation trend of corrected
CaO against SiO2 (Fig. 3), which is also influenced by
plagioclase accumulation, shows similar negative trends
with increasing SiO2 content. However, the corrected
Na2O contents are scattered and the differentiation trends
are less obvious.
Two distinct trends in the co-variation of FeO∗ (total
Fe as FeO) and MgO of Akagi volcano can be divided
into different rock series using the Miyashiro diagram
(Fig. 4a; Miyashiro, 1974). In this diagram, the Oe and
Om suites belong to the tholeiitic rock series, and the
Ol, Y and Cc suites to the calc-alkaline rock series. On
the other hand, although the Oe and Om suites are
slightly less depleted in FeO∗ than the Ol, Y and Cc
suites, the entire Akagi volcanic suite follows a calcalkaline differentiation trend in the AFM diagram (Fig.
4b). These major element characteristics of Akagi volcano
cannot be explained by a simple magmatic differentiation
process, because there appear to be considerable gaps in
the SiO2 and total alkali contents in the transition from
the Om to the Ol suites (Figs 2 and 4a). It may be,
therefore, necessary to invoke more complex processes,
such as magma mixing and crustal contamination associated with crystal fractionation. Moreover, the difference in the eruption style, which is the basis for the
discrimination of the stages, does not correspond to the
characteristics of the major element compositions. It may
thus be more suitable to discuss the geochemical evolution
of Akagi volcano without distinctions between stages and
substages. Here, we discriminate simply between samples
from the five stages (Oe, Om, Ol, Y, Cc), which are
distinguishable by geological discontinuities.
Trace element compositions
Trace element compositions of whole rocks are listed
in Table 1, and are plotted in primitive mantle
2307
AK1303
53·4
0·82
19·8
2·93
5·51
0·15
4·04
9·87
2·57
0·45
0·12
0·04
1·18
100·9
2·02
11·9
370
18·9
45·5
1·76
0·915
146
6·21
14·4
2·00
9·67
2·46
0·918
2·71
0·474
3·07
0·648
1·78
0·275
1·92
0·277
1·33
0·117
3·63
0·700
0·227
Sample no.:
wt %
SiO2
TiO2
Al2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K 2O
P 2O 5
H 2O−
H 2O+
Total
FeO∗/MgO
ppm
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
Oe
27·0
327
15·7
81·4
3·03
1·33
227
9·89
21·8
2·61
11·4
2·58
0·855
2·71
0·454
2·84
0·594
1·58
0·251
1·69
0·257
2·09
0·231
5·54
2·38
0·629
58·5
0·67
18·8
3·34
3·63
0·14
2·73
7·37
3·11
1·06
0·11
0·09
1·03
100·5
2·43
AK1302
17·8
357
15·8
63·1
2·18
0·892
185
7·57
17·4
2·24
9·85
2·40
0·813
2·46
0·416
2·58
0·558
1·48
0·236
1·59
0·237
1·86
0·150
4·27
1·69
0·381
56·3
0·66
19·1
3·38
3·88
0·14
3·10
7·87
2·73
0·66
0·13
0·24
2·10
100·3
2·23
AK2604
Om
24·3
379
23·7
107·7
4·04
0·539
280
12·83
24·6
3·29
14·2
3·22
1·01
3·31
0·579
3·26
0·703
1·90
0·290
2·02
0·298
2·58
0·257
4·53
2·43
0·507
57·1
0·79
19·2
2·70
4·39
0·13
2·73
7·67
3·10
0·95
0·18
0·17
1·16
100·3
2·49
AK1002
20·2
439
18·5
98·3
3·75
0·792
286
13·10
27·7
3·48
15·3
3·40
1·12
3·37
0·543
3·19
0·663
1·73
0·266
1·83
0·267
2·55
0·244
4·89
2·29
0·466
57·9
0·65
19·9
2·56
3·52
0·12
2·12
8·24
3·24
0·81
0·19
0·16
0·56
100·0
2·75
AK1004
23·1
436
19·2
92·3
3·50
0·713
280
11·99
25·6
3·31
14·6
3·06
1·04
3·07
0·489
2·93
0·610
1·55
0·239
1·66
0·245
2·31
0·223
5·38
2·14
0·483
58·2
0·64
19·9
2·25
3·55
0·11
1·94
7·95
3·24
0·86
0·20
0·20
0·86
99·8
2·88
AK1011
24·3
383
17·8
68·6
2·78
0·839
227
9·45
21·8
2·74
12·0
2·79
0·979
2·85
0·496
2·91
0·615
1·70
0·243
1·77
0·257
2·02
0·201
4·22
2·20
0·505
59·5
0·51
20·7
1·95
2·29
0·09
1·36
7·82
3·34
0·91
0·16
0·17
0·96
99·7
2·97
AK1103
19·5
340
16·3
64·7
2·06
0·494
180
8·04
17·7
2·33
10·3
2·52
0·868
2·57
0·414
2·74
0·556
1·47
0·225
1·55
0·232
1·63
0·142
4·10
1·66
0·374
53·6
0·82
19·1
3·58
5·84
0·16
4·53
8·55
2·52
0·43
0·14
0·12
0·75
100·2
2·00
AK0910
Ol
14·2
379
16·5
84·0
3·12
0·183
222
11·70
25·6
3·22
14·7
3·27
1·06
3·27
0·539
3·13
0·640
1·68
0·257
1·76
0·259
2·12
0·188
4·01
1·45
0·334
55·7
0·77
18·5
3·95
4·76
0·15
3·85
7·12
2·82
0·66
0·17
0·41
1·45
100·4
2·16
AK1010
Table 2: Major and trace element compositions of whole-rock samples from Akagi volcano
22·9
310
16·9
66·6
2·59
0·495
188
8·11
17·8
2·28
9·95
2·29
0·794
2·33
0·400
2·49
0·517
1·42
0·219
1·55
0·229
1·82
0·191
4·03
1·84
0·469
56·0
0·77
18·0
3·86
4·90
0·16
4·08
8·11
2·80
0·67
0·14
0·03
0·36
99·9
2·05
AK-A
22·0
325
15·6
70·7
3·17
0·475
201
8·67
18·8
2·37
10·4
2·46
0·830
2·61
0·429
2·64
0·537
1·42
0·211
1·44
0·217
1·99
0·245
6·63
2·06
0·476
56·4
0·77
18·0
4·61
3·78
0·15
4·05
7·72
2·82
0·80
0·13
0·07
0·71
100·0
1·96
AK1012
8·5
369
15·1
47·9
1·72
0·255
169
6·59
15·2
2·07
9·38
2·36
0·883
2·51
0·421
2·66
0·568
1·54
0·233
1·65
0·227
1·32
0·112
2·41
0·832
0·179
56·5
0·69
18·0
3·95
4·32
0·16
3·84
7·40
2·76
0·83
0·13
0·30
1·56
100·4
2·05
AK1201
22·0
383
25·5
95·1
4·08
1·08
334
16·7
31·4
4·39
19·0
4·03
1·22
4·18
0·684
4·12
0·876
2·26
0·358
2·46
0·372
2·56
0·247
4·85
2·46
0·522
59·5
0·62
17·6
4·29
2·69
0·13
2·70
6·49
2·83
0·75
0·15
0·48
1·68
99·9
2·42
AK1102
33·3
414
18·8
82·10
4·43
0·802
393
16·9
36·8
4·43
18·8
3·76
1·10
3·51
0·533
3·13
0·600
1·46
0·234
1·57
0·203
2·97
0·260
8·25
1·90
0·535
62·1
0·49
16·2
2·76
2·13
0·11
1·56
5·58
3·13
1·09
0·15
0·20
3·77
99·3
2·97
AK1108
JOURNAL OF PETROLOGY
VOLUME 42
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NUMBER 12
DECEMBER 2001
AK1202
58·0
0·69
18·0
3·29
3·54
0·13
3·06
6·96
2·78
0·86
0·12
0·62
1·90
100·0
2·12
22·4
316
17·4
101
3·17
1·19
351
13·6
28·8
3·34
13·8
2·93
0·941
3·05
0·477
2·96
0·637
1·69
0·266
1·87
0·274
2·71
0·194
6·29
3·15
0·704
Sample no.:
wt %
SiO2
TiO2
Al2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K 2O
P 2O 5
H 2O−
H 2O+
Total
FeO∗/MgO
ppm
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
Y
31·4
377
18·0
89·1
3·78
1·61
370
13·4
28·0
3·27
13·7
2·77
0·910
2·73
0·434
2·58
0·551
1·51
0·240
1·65
0·250
2·50
0·291
7·05
3·39
0·660
59·3
0·54
18·0
3·50
3·05
0·14
3·00
6·96
2·86
1·04
0·13
0·13
1·17
99·8
2·07
AK2601
31·0
391
14·1
79·3
3·37
1·61
351
12·6
27·4
3·03
12·2
2·46
0·840
2·34
0·379
2·22
0·480
1·27
0·201
1·43
0·215
2·23
0·246
7·47
3·46
0·627
60·3
0·57
17·1
3·25
3·35
0·14
3·03
6·59
2·98
1·10
0·14
0·03
0·97
99·6
2·07
AK2603B
30·9
369
17·2
93·1
3·66
1·68
422
13·6
28·6
3·25
13·5
2·70
0·859
2·62
0·415
2·43
0·535
1·45
0·222
1·62
0·232
2·51
0·264
7·20
3·79
0·617
60·5
0·56
17·8
3·23
2·78
0·12
2·44
6·84
3·06
1·12
0·14
0·03
0·81
99·4
2·33
AK2603G
31·9
366
17·8
89·5
3·51
1·55
356
13·5
27·7
3·30
13·5
2·88
0·894
2·63
0·428
2·59
0·545
1·47
0·238
1·64
0·249
2·29
0·240
5·45
3·41
0·659
60·6
0·53
17·6
2·73
3·30
0·13
2·72
6·82
3·04
1·08
0·13
0·08
0·99
99·8
2·12
32·2
386
17·8
92·9
3·80
1·58
377
14·6
29·4
3·57
14·5
2·97
0·957
2·88
0·470
2·63
0·602
1·62
0·256
1·78
0·262
2·46
0·289
6·86
3·38
0·654
61·1
0·55
17·1
3·33
2·92
0·13
2·95
6·63
2·86
1·10
0·12
0·09
0·94
99·8
2·00
AK2603W AK2602
30·1
341
13·8
92·7
2·84
1·31
353
12·7
27·5
3·12
12·4
2·54
0·892
2·52
0·380
2·39
0·483
1·36
0·213
1·48
0·219
2·50
0·193
5·97
3·44
0·692
62·0
0·55
17·5
2·61
2·73
0·11
2·36
6·42
3·20
1·17
0·12
0·18
0·83
99·8
2·15
AK2605
64·1
276
11·5
92·7
4·92
1·74
459
16·1
31·4
3·41
12·5
2·17
0·676
2·05
0·323
1·85
0·407
1·13
0·183
1·39
0·198
2·44
0·489
11·1
5·59
1·50
71·4
0·30
15·1
1·21
1·22
0·08
0·78
3·48
3·84
2·01
0·09
0·03
1·03
100·5
2·98
AK0807
Cc
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
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Fig. 2. Major element variation against SiO2 (wt %) content in Akagi samples. Dark and light gray shaded areas are the tholeiite and calcalkali rock series of the Nasu Volcanic zone in NE Japan, respectively [data from Kawano et al. (1961)].
(PM)-normalized diagrams in Fig. 5. The contents of
incompatible elements are enriched compared with most
basaltic rocks of the volcanic front of the NE Japan
arc (Shibata & Nakamura, 1997). However, the less
incompatible elements such as the HREE are less enriched. These trace element characteristics cannot be
explained by simple fractional crystallization from a
common primary basaltic magma in the volcanic front
using the phenocryst assemblage in the Akagi volcanic
rocks. Such a fractionation process should increase not
only the most incompatible trace elements but also the
HREE. The behavior of REE in the Akagi samples is
discussed below in more detail in a section describing
their isotopic compositions. With progress within a volcanic stage, the contents of highly incompatible elements
become gradually higher. However, no correlation is
observed between trace elements and SiO2 contents in
the Om and Ol stages. Akagi volcanic rocks show positive
spikes of Sr and Pb, and remarkably negative spikes for
Nb and Ta, consistent with the general features of trace
elements in island-arc volcanic rocks (e.g. Wood et al.,
1979; Perfit et al., 1979; Sakuyama & Nesbitt, 1986;
Shibata & Nakamura, 1997; Woodhead et al., 1998).
Sr–Nd–Pb isotope systematics of Akagi
volcano
The whole-rock Sr, Nd and Pb isotopic compositions are
listed in Table 3 and are plotted in Fig. 6 along with the
reference variations of volcanic rocks from the NE Japan
arc, lower-crustal materials, oceanic sediments and
2310
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Fig. 3. Corrected Al2O3, CaO and Na2O contents plotted against SiO2
content by the subtraction of phenocryst compositions from the bulk
Akagi samples. Symbols are the same as in Fig. 2. The continuous and
broken lines indicate corrected compositions of the tholeiite and calcalkali rock series of Adatara volcano, respectively (Fujinawa, 1988,
1990).
MORBs. The 87Sr/86Sr ratios of the Akagi samples vary
widely from 0·70603 to 0·70879, consistent with the data
reported by Notsu et al. (1985), and are significantly higher
than those of other typical volcanoes in the northern part
of Tanakura Tectonic Line in the volcanic front of the
NE Japan arc, which have an average 87Sr/86Sr ratio of
>0·7045 (Notsu, 1983; Gust et al., 1997; Shibata &
Nakamura, 1997). The Nd isotopic compositions of Akagi
range from −0·4 to −8·6 in Nd, and are much lower
than those of other volcanoes in the northern part of the
volcanic front with Nd of 3–10 (Gust et al., 1997; Shibata
& Nakamura, 1997). The 206Pb/204Pb, 207Pb/204Pb and
208
Pb/204Pb ratios vary in the range of 18·21–18·44,
15·56–15·64 and 38·44–38·79, respectively (Fig. 6b and
c). These isotope characteristics of Akagi, therefore, are
regarded as ‘unusual’ in the NE Japan arc, confirming the
original observations based on Sr isotopic compositions by
Notsu et al. (1985).
Fig. 4. (a) FeO∗/MgO plotted against SiO2 for Akagi volcano. The
TH series and CA series were defined by Miyashiro (1974), and the
boundary line is defined by FeO∗/MgO = 0·1562 × SiO2 − 6·685.
(b) AFM diagram for Akagi volcano. A, Na2O + K2O; F, FeO +
0·9Fe2O3; M, MgO. Symbols are as in (a). The boundary line between
the TH and CA series is based on the definition of Gill (1981).
In the Sr and Nd isotope diagram (Fig. 6a), the Akagi
samples have a linear trend with an extremely isotopically
enriched signature compared with those of other ‘typical’
volcanoes in the volcanic front of the NE Japan arc,
which have simple isotopic across-arc variations (Shibata
& Nakamura, 1997). These isotopic compositions are
widely varied not only over the entire history of the
volcano but also with each volcanic stage of Akagi
volcano. An extrapolation of the Akagi trend to the
depleted direction nearly intersects the isotopic trend of
the NE Japan arc at a value typical of volcanoes on the
front. As is shown in Fig. 6b and c, the Pb isotopic
compositions of the Akagi samples form clusters that are
clearly distinct from the isotopic trend of the NE Japan
arc (Shibata & Nakamura, 1997). The 206Pb/204Pb ratios
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DECEMBER 2001
Fig. 5. Primitive mantle (PM)-normalized trace element patterns for whole-rock samples from Akagi volcano. The PM normalization values
are from McDonough & Sun (1995) in this and subsequent figures. The shaded area shows a range of samples from the Iwate and Funagata
volcanoes in the volcanic front of the NE Japan arc (Shibata & Nakamura, 1997).
are lower, whereas the 207Pb/204Pb and 208Pb/204Pb ratios
are higher than those of the ‘typical’ volcanoes on the
front. In the Nd–206Pb/204Pb diagram (Fig. 6d), the Akagi
data also define a linear trend clearly different from
the tendency in the isotope variations of other primary
magmas in the northern part of the NE Japan arc, and
which extrapolates to the values characteristic of the
volcanic front.
The above isotope systematics suggest that the Akagi
volcanic rocks were formed by two-component mixing
of an isotopically depleted and an enriched end-member.
Intersections between the ‘typical’ trend of the NE Japan
arc and the extrapolations of the Akagi trends in the
Sr–Nd and Nd–Pb diagrams indicate that the depleted
end-member is similar to the primary magma in the
volcanic front of the NE Japan arc, which was defined
by Shibata & Nakamura (1997). It is, however, still
difficult to identify the source of the enriched end-member
using these isotope systematics.
On the basis of the existence of an enriched component
deduced from the isotope systematics and the tectonic
setting of the Akagi volcano, the following candidates for
the enriched end-member can be postulated: (1) oceanic
sediments; (2) upper crust; (3) lower crust. Isotopic variations of oceanic sediments in the NW Pacific and
lower-crustal materials are shown in Fig. 6. The isotopic
composition of JG-1, the granitic standard of the Geological Survey of Japan, which outcrops near Akagi
volcano and is considered to be the upper-crustal basement of Akagi, is shown in the diagrams. It should be
noted that the JG-1 composition lies within the field for
oceanic sediments (data from Nohda & Wasserberg,
1981; Koide & Nakamura, 1990), therefore the uppercrustal field is not shown in Fig. 6. The isotopic characteristics of lower-crust materials under the NE Japan
arc have not been well understood. The lower crust
under the arc might be considered as a fragment of
continental material and/or a piled-up sequence of young
2312
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Table 3: Sr, Nd and Pb isotopic compositions of whole-rock samples from Akagi volcano
Sample no.
87
Sr/86Sr
143
Nd/144Nd
Nd
206
Pb/204Pb
207
Pb/204Pb
208
Pb/204Pb
Cc
0·706852±9
0·512437±12
−3·92
18·444±9
15·600±7
38·667±18
AK2605
0·707414±10
0·512297±11
−6·65
18·290±7
15·573±6
38·599±15
AK2602
0·708087±11
0·512248±12
−7·61
18·268±35
15·592±29
38·695±73
AK2603W
0·707774±10
0·512274±12
−7·10
18·280±13
15·578±11
38·630±27
AK2603G
0·707774±9
0·512233±10
−7·90
18·285±24
15·583±21
38·645±52
AK2603B
0·707800±11
0·512296±9
−6·67
18·294±23
15·594±20
38·677±49
AK2601
0·708114±11
0·512209±13
−8·37
18·290±64
15·609±55
38·732±135
AK1202
0·707385±10
0·512336±12
−5·89
18·309±12
15·577±10
38·605±24
AK1108
0·708792±12
0·512199±13
−8·56
18·262±20
15·583±16
38·625±42
AK1102
0·708244±10
0·512210±10
−8·35
18·259±19
15·584±18
38·634±41
AK1201
0·706833±8
0·512442±12
−3·82
18·366±20
15·576±4
38·580±9
AK0807
Y
Ol
AK1012
0·707119±8
0·512392±10
−4·80
18·327±29
15·576±24
38·597±60
AK-A
0·706765±8
0·512423±11
−4·19
18·341±79
15·602±65
38·642±165
AK1010
0·708632±9
0·512224±12
−8·08
18·242±74
15·605±61
38·684±156
AK0910
0·706857±9
0·512430±10
−4·06
18·279±75
15·575±65
38·556±157
AK1103
0·706840±9
0·512440±9
−3·86
18·380±50
15·637±39
38·794±105
AK1011
0·708705±9
0·512213±10
−8·29
18·206±22
15·583±19
38·643±47
AK1004
0·708716±10
0·512210±8
−8·35
18·253±16
15·625±91
38·734±156
AK1002
0·707532±8
0·512381±10
−5·01
18·302±20
15·579±16
38·615±41
AK2604
0·706972±10
0·512388±12
−4·88
18·334±97
15·614±81
38·665±204
AK1302
0·707207±8
0·512387±10
−4·90
18·375±45
15·600±38
38·556±95
AK1303
0·706034±10
0·512618±11
−0·39
18·354±29
15·561±23
38·436±60
Om
Oe
Isotopic fractionation was normalized to 86Sr/88Sr = 0·1194,
2 mean.
Nd/144Nd = 0·7219. Analytical precisions for isotope data are
146
metamorphic rocks originally derived from the arc
magmas (Arculus & Johnson, 1981). In the latter case,
the isotopic characteristics should be similar to those of
the arc volcanic rocks, as a result of less time-integrated
isotopic evolution. We prefer, therefore, that the isotopic
characteristics of the lower-crustal materials are similar
to ‘continental’ type lower crust in Fig. 6.
In the Sr–Nd diagram (Fig. 6a), the isotopic compositions of oceanic sediments and lower crust overlap,
making it difficult to unambiguously identify an enriched
end-member. However, the Pb–Pb and Nd–Pb diagrams
clearly discriminate between lower crust and oceanic
sediments and/or upper crust. On the basis of Fig. 6b–d,
it appears that oceanic sediments and/or upper crust are
not possible candidates for the enriched end-member. In
practice, the isotopic compositions of sediments, which
occur on the Philippine Sea plate, cannot be the enriched
end-member for Akagi volcano. Consequently, it is most
likely that the enriched end-member involved in magma
formation of Akagi volcano is lower-crustal material,
although Notsu et al. (1985) proposed a sediment component derived from the subducting Philippine Sea plate
as the enriched end-member in Akagi volcano based on
87
Sr/86Sr and 18O data. The complete Sr, Nd and Pb
isotopic dataset from the along-arc volcanoes by Kersting
et al. (1996) and Gust et al. (1997) is also shown in Fig.
6. The isotopic trends defined by the volcanoes located
to the south of the Tanakura Tectonic Line (STTL:
Akagi, Nikkoshirane, Nantai, Takahara, Nasu) essentially
have the same direction as those of the primary magmas
in the volcanic front of the NE Japan arc (Fig. 6).
Therefore, it is likely that the depleted end-member for
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DECEMBER 2001
Fig. 6. Nd–Sr–Pb isotope systematics of the Akagi samples. Symbols are as in Fig. 2. Nd isotope ratios are normalized to CHUR ( 143Nd/
144
Nd = 0·512638). The large open circle represents the isotopic composition of the inferred primary magma of the NE Japan arc, and the
thick continuous line extending from the open circle indicates the across-arc variation in the NE Japan arc (Shibata & Nakamura, 1991, 1997).
The fields ‘N’ and ‘S’ represent the along-arc variations of the magmatic suites from north and south of the Tanakura Tectonic Line (NTTL
and STTL), respectively (Kersting et al., 1996; Gust et al., 1997). ×, isotopic composition of JG-1, the granitic rock standard material of the
Geological Survey of Japan (Nohda & Wasserburg, 1981; Koide & Nakamura, 1990). The range of isotopic variation of MORB, oceanic
sediments including those on the Philippine Sea plate and the lower crust are also shown. Data source for MORB: Cohen et al. (1980), Dupre
& Allègre (1980), White & Hofmann (1982), Le Roex et al. (1983); oceanic sediments: O’Nions et al. (1978), White & Patchett (1984), Woodhead
& Fraser (1985), White & Dupre (1986), Ben Othman et al. (1989), Cousens et al. (1994), Shimoda et al. (1998); lower crust: Zartman & Doe
(1981), Rudnick et al. (1986), Zartman & Haines (1988), Stoltz & Davies (1989), Kempton et al. (1990), Rudnick & Goldstein (1990), Downes et
al. (1991).
the source of the volcanoes from the STTL also is similar
to the MORB-type wedge mantle metasomatized by the
fluid derived from dehydration of the subducted slab as
proposed by Shibata & Nakamura (1997) for the primary
magma source of the NE Japan arc. Moreover, the
Nd–Pb diagram (Fig. 6d) for the STTL group clearly
indicates the involvement of lower-crustal materials as
an enriched end-member, as well as at Akagi volcano.
It is, therefore, not necessary to introduce a unique
mantle source such as Indian Ocean-type mantle beneath
the volcanoes from the STTL to explain the isotope
systematics of the STTL group based on the dataset
given in this study and in previous studies (Kersting et
al., 1996; Gust et al., 1997).
Major and trace element compositions of
phenocrysts
Seven samples from the Ol stage, which is characterized
by large variations in whole-rock major element and
isotopic composition (SiO2 = 55·8–63·5 wt %, 87Sr/
86
Sr = 0·70683–0·70879), were selected to investigate
detailed magmatic processes based on the chemical and
isotopic compositions of phenocrysts.
Phenocrysts in Akagi volcano rocks are always zoned.
The mg-numbers [defined as 100 × Mg/(Mg + Fe2+)]
of cores in orthopyroxene and clinopyroxene are 75–65
and 80–73, respectively. Reverse zoning is commonly
observed in the pyroxene of samples AK1102, AK0910,
AK1010 and AK1108. The an-number in the cores of
2314
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
plagioclase phenocrysts does not have a bimodal signature, which has been regarded as evidence of magma
mixing according to Sakuyama (1981). Hornblende
phenocrysts occur in only the most siliceous sample in
this stage. The hornblende phenocrysts have a zonal
structure with decreasing mg-number and edenite component from core to rim, and a pseudomorph of hornblende phenocrysts, consisting of fine-grained plagioclase
+ orthopyroxene + clinopyroxene + opaque minerals,
also occurs in five siliceous samples listed in Table 1.
Trace element compositions of orthopyroxene, clinopyroxene and plagioclase phenocrysts in the Ol stage are
given in Table 4 and their primitive mantle-normalized
patterns are shown in Figs 7 and 8. The trace element
patterns of pyroxenes show clear negative Sr and Nb
spikes. Experimental determinations of the Kd for Sr
between pyroxenes and silicate melt show no Sr depletion
compared with the adjacent elements Pr and Nd (Green,
1994). The Sr depletion, therefore, indicates that the
pyroxenes crystallized from Sr-depleted melts, although
Sr enrichment is observed in whole-rock analyses of the
Ol samples (see Fig. 5). Such Sr enrichment could be
expected as a result of the accumulation of plagioclase,
which is characterized by a strong Sr positive spike (Fig.
8). The influence of plagioclase accumulation is expected
based on the major element compositions of the wholerock samples (see Fig. 3). The positive Sr spikes are
removed by consideration of phenocryst subtraction; for
example, Sr/Sr∗ value [Sr∗ is defined as (Pr + Nd)/2
on the primitive mantle-normalized value] of the melt
part in AK-A decreases from 2·0 to 0·7.
The Nb depletions in the phenocrysts can be attributed
to the whole-rock compositions that are unique to islandarc volcanic rocks (see also Fig. 5). The Zr depletion
in clinopyroxene and the slightly positive Zr spike in
orthopyroxene may be simply explained by the difference
in Kd [i.e. KdHFSE/KdLILE,LREE >1 for orthopyroxene, KdHFSE/
KdLILE,LREE <1 for clinopyroxene (where HFSE indicates
high field strength elements, LILE indicates large ion
lithophile elements and LREE are light REE); Green,
1994]. The Y depletion in the plagioclase phenocrysts
may be a characteristic feature of plagioclase, because
such depletion is always observed in the plagioclases used
for our secondary ion mass spectrometry standards (also
precisely characterized by ICP-MS). In samples AK1010
and AK1108, the incompatible elements in the orthopyroxene rims are significantly depleted compared with
those of the core. This is consistent with reverse zoning
(based on mg-number) in the pyroxene of these samples.
Sr and Nd isotope compositions of
phenocrysts in the Ol stage
The isotopic compositions of Sr and Nd for phenocrysts
in the Ol stage are given in Table 5 and plotted in a
Sr–Nd isotope diagram in Fig. 9. In this diagram, the
isotopic compositions of the melts are calculated using
the Sr and Nd isotopic and elemental compositions of
the whole rocks, and the modal abundance, densities,
isotopic and elemental compositions of the phenocrysts.
The results are given in Table 6. The isotopic compositions of the phenocrysts tend to have more enriched
compositions, especially for Sr, exceeding the analytical
uncertainties, than the calculated compositions of the
melts. In other words, the phenocrysts are isotopically in
disequilibrium with the surrounding melt.
The uncertainty of calculated melt isotopic compositions mainly depends on the measured modal abundance of the phenocrysts, especially plagioclase, which is
a main sink for Sr. The statistical error of the plagioclase
modal abundance is estimated to be >3%, based on the
counting statistics. This uncertainty affects the 87Sr/86Sr
ratio of melts up to ±>0·0002 for AK-A and AK1012,
but the uncertainties in the others are less than ±0·0001.
In contrast, the calculated Nd isotopic compositions of
the melts are not significantly affected by the uncertainty
in the modal abundance. Heterogeneity in the sample
also causes the error in the estimation of the Sr and Nd
isotopic composition of the melt. However, the distributions of phenocrysts are homogeneous in the samples
from the Ol stage. Therefore, the isotopically depleted
characteristics of the calculated melts compared with the
phenocrysts are apparently significant except in sample
AK1010. In AK1010, the Sr isotopic composition of the
melt is identical to that of the phenocrysts within the
analytical uncertainties. However, the Nd isotopic composition of the plagioclase is significantly lower than that
of the calculated melt.
INTERACTION BETWEEN ISLANDARC MAGMA AND LOWER CRUST
In Fig. 10, chondrite-normalized REE patterns of Akagi
volcano samples are compared with those for Adatara
volcano (Fujinawa, 1992), which is a neighboring volcano
in the volcanic front of the NE Japan arc and has similar
petrological and mineralogical features to those of Akagi.
This comparison shows that Akagi samples have (1)
significantly lower HREE abundance, (2) significantly
higher LREE/HREE ratios, and (3) U-shaped REE
patterns between the middle REE (MREE) and HREE.
The REE patterns from Adatara volcano were explained
by crystal fractionation (Fujinawa, 1992). However, the
REE characteristics of Akagi volcano cannot be formed
by such a simple fractionation involving removal of
minerals such as plagioclase, orthopyroxene, clinopyroxene and opaque minerals occurring as phenocrysts
in the Akagi samples. Because the mineral–melt partition
coefficients of the REE are much smaller than one,
2315
2316
43·4
25·6
12·5
0·002
27·1
1·00
3·93
0·825
5·27
2·37
0·935
3·44
3·81
2·03
2·16
0·317
ppm
Sr
Y
Zr
Nb
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Dy
Er
Yb
Lu
21·3
37·3
23·3
0·001
0·492
1·93
7·78
1·76
10·8
5·06
1·21
5·90
7·50
3·29
4·00
0·516
20·7
34·1
21·4
0·002
0·284
1·73
7·13
1·89
11·6
4·74
1·26
5·80
7·22
3·21
3·52
0·470
core
0·207
4·15
1·54
0·001
0·225
0·018
0·045
0·013
0·109
0·081
0·027
0·140
0·479
0·348
0·511
0·074
0·291
3·33
1·18
0·001
0·489
0·030
0·072
0·010
0·100
0·068
0·028
0·163
0·297
0·276
0·449
0·075
core
rim
0·204
5·89
1·54
n.d.
0·016
0·014
0·078
0·023
0·170
0·090
0·046
0·207
0·697
0·416
0·642
0·121
rim
0·557
5·46
2·98
0·001
0·687
0·037
0·074
0·025
0·157
0·137
0·046
0·208
0·610
0·476
0·587
0·094
orthopyroxene
16·7
21·4
10·8
0·001
2·63
1·08
3·42
0·711
4·73
2·30
0·579
2·71
3·15
1·68
1·91
0·274
525
0·054
0·490
0·193
12·1
0·903
1·70
0·197
0·948
0·191
0·430
0·222
0·085
0·031
0·033
0·004
rim
plagioclase
558
0·026
0·032
0·065
8·81
0·639
1·14
0·156
0·647
0·095
0·254
0·038
0·032
0·019
0·008
0·003
rim
plagioclase
596
0·031
0·091
0·032
11·7
1·75
2·88
0·365
1·16
0·189
0·539
0·087
0·080
0·033
0·006
0·002
core
621
0·056
0·324
0·152
11·8
2·28
4·12
0·433
1·69
0·276
0·496
0·219
0·054
0·028
0·015
0·003
core
25·6
31·7
29·0
0·002
0·106
1·71
7·72
1·53
10·8
4·19
1·20
5·27
5·51
2·93
2·79
0·392
core
22·0
33·9
22·3
0·002
0·403
1·48
6·01
1·44
10·1
3·98
0·960
5·39
5·54
2·91
2·91
0·445
rim
19·2
34·1
23·2
0·002
2·66
1·49
6·47
1·65
10·1
4·25
1·04
5·87
6·80
3·55
3·21
0·479
core
clinopyroxene
AK1012
28·4
25·3
21·6
0·005
0·96
1·40
6·22
1·40
9·53
3·05
1·00
4·24
3·81
2·20
2·05
0·289
rim
clinopyroxene
AK1010
0·889
3·46
2·02
0·001
1·20
0·037
0·091
0·028
0·112
0·080
0·025
0·196
0·415
0·327
0·421
0·077
core
0·226
4·82
1·92
0·001
0·108
0·012
0·063
0·013
0·089
0·052
0·062
0·154
0·478
0·485
0·856
0·134
rim
0·205
4·10
1·51
0·001
0·231
0·011
0·064
0·015
0·114
0·046
0·047
0·198
0·492
0·394
0·576
0·089
core
orthopyroxene
0·207
2·62
1·49
0·001
0·037
0·012
0·026
0·009
0·052
0·109
0·029
0·162
0·305
0·247
0·300
0·041
rim
orthopyroxene
596
0·043
0·194
0·137
17·0
1·75
3·43
0·360
1·44
0·287
0·526
0·287
0·082
0·033
0·026
0·004
rim
plagioclase
667
0·051
0·355
0·0821
38·5
2·78
5·00
0·621
1·85
0·333
1·24
0·317
0·186
0·068
0·037
0·004
rim
plagioclase
583
0·097
2·77
0·607
24·9
2·59
4·35
0·487
1·87
0·213
0·710
0·261
0·125
0·033
0·017
0·004
core
655
0·091
0·282
0·181
11·8
1·46
2·78
0·367
1·50
0·211
0·504
0·183
0·107
0·026
0·025
0·002
core
VOLUME 42
clinopyroxene
Sample no.: AK-A
ppm
Sr
Y
Zr
Nb
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Dy
Er
Yb
Lu
core
rim
rim
core
orthopyroxene
clinopyroxene
Sample no.: AK0910
Table 4: Trace element compositions of phenocrysts from Ol stage samples obtained by ion microprobe
JOURNAL OF PETROLOGY
NUMBER 12
DECEMBER 2001
2317
ppm
Sr
Y
Zr
Nb
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Dy
Er
Yb
Lu
61·4
188·8
81·8
0·335
81·4
8·33
37·0
8·90
58·0
20·5
4·39
32·1
29·5
15·4
16·0
2·56
58·4
191·8
81·3
0·312
72·8
7·49
32·8
7·83
53·1
18·4
4·24
28·1
29·4
14·8
15·9
2·31
0·154
5·72
3·75
0·001
0·071
0·011
0·049
0·013
0·139
0·125
0·035
0·143
0·611
0·573
0·811
0·140
0·176
5·86
2·83
0·001
0·102
0·101
0·297
0·042
0·250
0·174
0·041
0·370
0·602
0·599
0·918
0·137
core
rim
core
0·447
6·63
2·74
0·001
0·928
0·022
0·090
0·012
0·131
0·118
0·047
0·286
0·720
0·552
0·852
0·115
rim
0·162
6·47
3·48
0·001
0·106
0·008
0·071
0·015
0·173
0·099
0·037
0·322
0·727
0·513
0·776
0·115
orthopyroxene
AK1108
Sample:
21·1
27·9
14·6
0·002
0·046
1·09
5·09
1·30
7·89
3·28
0·972
4·20
5·92
2·51
2·57
0·424
hornblende
22·1
32·6
19·4
0·002
0·231
1·48
6·14
1·45
9·62
3·75
1·05
6·01
6·34
3·28
3·07
0·478
ppm
Sr
Y
Zr
Nb
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Dy
Er
Yb
Lu
core
rim
rim
core
orthopyroxene
clinopyroxene
Sample no.: AK1201
764
0·040
0·351
0·114
17·6
3·67
5·96
0·645
2·23
0·269
0·933
0·172
0·069
0·041
0·010
0·003
rim
plagioclase
612
0·043
0·840
0·043
10·4
1·48
2·66
0·301
1·00
0·186
0·492
0·253
0·079
0·026
0·007
0·001
rim
plagioclase
800
0·036
0·082
0·070
10·8
2·94
4·49
0·560
1·61
0·300
0·638
0·292
0·045
0·030
0·006
0·002
core
527
0·030
0·012
0·022
2·49
0·391
0·791
0·082
0·375
0·099
0·180
0·105
0·025
0·010
0·006
n.d.
core
26·1
46·8
41·9
0·008
22·4
2·61
10·6
2·42
16·3
5·92
1·39
7·47
8·23
3·92
3·69
0·643
rim
17·0
42·7
25·1
0·002
4·13
1·89
8·53
2·10
13·3
4·58
1·47
5·48
8·73
4·02
4·21
0·583
core
clinopyroxene
AK1102
0·177
5·67
2·33
0·001
0·097
0·020
0·107
0·027
0·164
0·073
0·017
0·177
0·454
0·658
0·783
0·131
rim
0·140
5·09
2·71
0·001
0·073
0·005
0·032
0·010
0·084
0·032
0·034
0·167
0·370
0·477
0·752
0·110
core
orthopyroxene
681
0·052
0·153
0·057
12·9
2·16
3·88
0·437
1·62
0·248
0·644
0·323
0·055
0·040
0·017
0·003
rim
plagioclase
774
0·034
0·046
n.d.
12·0
3·37
5·46
0·588
1·84
0·274
0·752
0·029
0·056
0·025
0·014
0·002
core
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
DECEMBER 2001
Fig. 7. PM-normalized trace element patterns of clinopyroxene and orthopyroxene phenocrysts in Ol stage samples. For the AK1108 sample,
hornblende data are represented instead of clinopyroxene, because there are no clinopyroxene phenocrysts in this sample. Open and filled
symbols indicate the rim and core of phenocrysts, respectively.
fractional crystallization in a shallow magma chamber
would increase the total REE with LREE enrichment,
as is observed in Adatara volcano. This mechanism,
therefore, cannot explain characteristics (1) and (2) of
Akagi’s REE compositions.
On the basis of a high-pressure melting experiment
involving a basaltic composition, Forden & Green
(1992) suggested that the liquidus phase is clinopyroxene
under high-H2O conditions in the lower-crustal environment, and olivine is the liquidus phase at lower
pressure conditions. When the temperature of the melts
drops, clinopyroxene becomes unstable, and reacts with
the melt to form amphibole, resulting in the evolved
melt being more siliceous and decreasing the H2O
content of the magma itself. The lower-crustal granulites
commonly contain clinopyroxene; ascending H2O-rich
magma from the mantle, therefore, might react with
the clinopyroxene and fractionate amphibole as a
residual phase. The MREE are preferentially partitioned
into amphibole, with mineral–melt partition coefficients
greater than one (Adam & Green, 1994; Sisson, 1994;
Witt-Eickschen & Harte, 1994). Consequently, the
evolved magma acquires a U-shaped REE pattern
because of the buffering effect of residual amphibole.
Partial melting experiments on lower-crustal materials
under high-pressure conditions indicate that the resultant liquid has a felsic composition caused by the
persistence of mafic minerals in the residue (e.g. Beard
& Lofgren, 1991; Beard et al., 1993, 1994). Although
amphibole fractionation processes in shallow magma
chambers could also explain the U-shaped REE patterns, the occurrence of hornblende phenocrysts is rare
and only very small amounts occur in felsic samples
in the later volcanic stages (Table 1). This process is,
therefore, unlikely to explain the overall U-shaped
signature retained in the Akagi volcanic rocks.
If the interaction between H2O-rich magma and lowercrustal material is a major process controlling the REE
patterns of Akagi volcano, a process of ‘assimilation–
fractional crystallization’ (AFC, DePaolo, 1981) under
2318
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Fig. 8. PM-normalized trace element patterns of plagioclase phenocrysts in Ol stage samples. Open and filled symbols indicate the rim and
core of phenocrysts, respectively.
lower-crustal conditions might be invoked. To examine
the assimilation process in the lower crust, the AFC
model (DePaolo, 1981) was applied to the isotope and
trace element systematics in the Akagi volcano. In this
model calculation, a primary magma deduced from ‘normal’ primary magma on the northern part of the volcanic
front in the NE Japan arc, defined by Shibata & Nakamura (1997), is assumed as an ‘assimilator’. It is, however, difficult to determine the other parameters such
as isotope and element ratios of assimilated material
(hereafter denoted as ‘assimilant’) under Akagi volcano,
and bulk partition coefficients between the melt and the
assimilant. Thus, we fixed these parameters, and only
evaluated the assimilation rate (r) as given by
r = Ma(t)/Mc(t) (DePaolo, 1981)
where Ma(t) and Mc(t) represent a rate (mass/unit time)
of wallrock assimilation and a rate of fractionation of
crystallizing phases, respectively. Thus, the assimilation
rate implies the rate of contribution of masses controlled
by the assimilation and the crystal fractionation to the
total mass of magma. Furthermore, the assimilation rate
is an essential parameter to evaluate the temperature
conditions at which the AFC takes place. The averages
of element ratios of lower-crustal materials compiled by
Rudnick & Fountain (1995) are used for the assimilant.
The isotopic compositions of the assimilant are obtained
by the extrapolation of the linear correlations in the
multi-isotope systematics of Akagi volcano. The bulk
partition coefficients between magma and the assimilant
were obtained by using the modal abundance of residual
phases consisting of plagioclase, amphibole, clinopyroxene and orthopyroxene, as determined by hydrous
melting experiments at lower-crustal conditions by Beard
& Lofgren (1991). The REE signature of Akagi volcanic
rocks indicates that amphibole fractionation was involved
in the evolution of magma under lower-crustal conditions
as discussed in the previous sections. All parameters used
in the AFC calculations are compiled in Table 7.
2319
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
DECEMBER 2001
Table 5: Isotopic analyses of phenocysts in Ol
stage samples
87
Sample no.
AK1108
AK1102
AK1201
AK1012
AK-A
AK1010
AK0910
Sr/86Sr
143
Nd/144Nd
Nd
Pl
0·709051±10
0·512218±7
−8·19
Opx
0·708974±11
0·512195±7
−8·64
Hb
0·708919±10
0·512143±7
−9·66
Pl
0·708570±10
0·512191±16
−8·72
Opx
0·708228±9
0·512216±18
−8·23
Cpx
0·708221±7
0·512212±17
−8·31
Pl
0·707046±12
0·512422±12
−4·21
Opx
0·707267±13
0·512455±30
−3·57
Cpx
0·707310±10
0·512364±6
−5·34
Pl
0·707596±11
0·512401±9
−4·62
Opx
0·707275±13
0·512381±11
−5·01
Cpx
0·707375±9
0·512401±32
−4·62
Pl
0·707200±10
0·512371±10
−5·21
Opx
0·706833±9
0·512420±31
−4·25
Cpx
0·707245±11
0·512384±6
−4·95
Pl
0·708631±10
0·512139±12
−9·73
Opx
0·708790±20
0·512234±21
−7·88
Cpx
0·708786±26
0·512277±11
−7·04
Pl
0·707032±11
0·512434±11
−3·98
Opx
0·706928±9
0·512429±10
−4·08
Cpx
0·707086±39
0·512445±12
−3·76
Pl, plagioclase; Opx, orthopyroxene; Cpx, clinopyroxene; Hb,
hornblende. Isotopic fractionation was normalized to 86Sr/
88
Sr = 0·1194, 146Nd/144Nd = 0·7219. Analytical precisions for
isotope data are 2 mean.
Representative results from the AFC model calculations
are presented together with data for Akagi, Adatara and
other volcanoes from NE Japan in Fig. 11a–d. There
appear to be correlations between isotope and element
ratios in the arc data, although the correlation is poor
for Pb isotopes, probably as a result of the small variation
of Pb isotope compositions and the small differences
between those of the lower crust and the primary magma.
These correlations are obviously different from those of
basaltic rocks elsewhere in the NE Japan arc and at
Adatara volcano, for which the source heterogeneity and
crystal fractionation processes have caused variations in
isotope and elemental ratios without crustal assimilation
(Shibata & Nakamura 1997). As shown in Fig. 11, most
of the Akagi data plot near the mixing lines with an
assimilation rate of 0·9, although those of the earliest
and final stages tend to plot above the mixing lines with
assimilation rates <0·9. This observation indicates that
the extent of crustal assimilation relative to the crystallization in the middle stage is larger than in the earliest
Fig. 9. Sr and Nd isotope systematics of phenocryst minerals in Ol
stage samples. Data from the same sample are connected by tie-lines.
The bulk-rock and calculated melt isotopic compositions (see text) are
also shown.
and latest stages of Akagi volcano. In the middle stage,
therefore, the enrichment of incompatible elements together with isotope variations may largely have resulted
from significant lower-crustal assimilation by primary
magma similar to ‘normal’ basaltic magmas in the northern part of the volcanic front in the NE Japan arc.
2320
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Table 6: Calculated Sr and Nd compositions,
and isotope compositions of melt part in Ol
stage samples
Sample no. Sr (ppm)
87
AK1108
323
AK1102
231
AK1201
Sr/86Sr
Nd/144Nd Nd
Nd (ppm)
143
0·70859
24·5
0·51220
−8·5
0·70770
29·5
0·51221
−8·3
360
0·70667
15·8
0·51244
−3·8
AK1012
246
0·70632
19·2
0·51239
−4·8
AK-A
199
0·70591
17·5
0·51243
−4·1
AK1010
327
0·70863
19·7
0·51223
−8·0
AK0910
278
0·70666
16·8
0·51243
−4·1
Phenocryst mode as in Table 1. Sr and Nd isotope compositions of whole rocks and phenocrysts as in Tables 3 and
5. Trace element compositions of whole rocks and phenocrysts as in Tables 2 and 4. Densities (in g/cm3): melt, 2·6; pl,
2·6; cpx and opx, 3·4; hb,3·0.
When the assimilation rate is nearly 1·0, the latent
heat of crystallization and the heat of fusion for the
assimilation process are thermally balanced, approximately. Such a circumstance is likely to be achieved
under lower-crustal conditions, because the geotherm in
the lower crust is such that temperatures are 600–700°C,
making it feasible to reach the solidus temperature of the
assimilant by intrusion of magma and with less heat loss.
On the other hand, the latent heat of crystallization could
be mostly consumed by raising the temperature of the
wallrock to its solidus, when assimilation occurs under
shallower crustal conditions. Consequently, this results
in an assimilation rate smaller than that occurring under
lower-crustal conditions. From the above discussion based
on the AFC model calculations, it may be concluded
that the assimilation occurred primarily in the lowercrustal region, resulting in the unique isotope characteristics of the magmas of Akagi volcano.
The seismologically determined Moho depth gradually
increases from the northern part of NE Japan along the
volcanic front, and it shows the maximum value beneath
Akagi volcano (>38 km, >1·3 GPa), which is >4 km
deeper than the typical depth beneath the volcanic front
of the NE Japan arc (Zhao et al., 1992). Hildreth &
Moorbath (1988) suggested that chemical interaction
between magma and crust is closely related to the crustal
thickness in the Andean arc. The thicker crust might
effectively increase the assimilation rate, because the
higher temperatures in the bottom of the crust mean
that it more easily attains its solidus temperature. Consequently, volcanoes such as Akagi located on thicker
crust might experience a larger effect of assimilation of
lower crust on their trace element and isotope compositions.
EXISTENCE OF A WATER-RICH
MAGMA
The above discussion based on isotope and trace element
systematics requires a primary magma sufficiently enriched in H2O to stabilize amphibole in the lower crust.
It is, therefore, essential to understand the mechanism
whereby an H2O-rich source region for the primary
magma in the mantle wedge beneath Akagi volcano
could form.
Fig. 10. Chondrite-normalized REE patterns for the TH and CA series of Akagi samples. The normalizing values are from McDonough &
Sun (1995). The shaded area shows the range of REE patterns for the TH and CA series samples from Adatara volcano (Fujinawa, 1992).
2321
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
DECEMBER 2001
Table 7: The parameters used for model calculation and the lower-crustal assimilation
Element
Plag. Kd∗
Cpx Kd†
Opx Kd†
Amph. Kd‡
Whole-rock
Calc. melt
composition§
composition
(ppm)
(ppm)
La
0·15
0·04
0·008
0·16
8
35
Sr
3·1
0·12
0·018
0·50
348
297
11
Nd
0·12
0·19
0·030
0·77
Pb
0·39
0·10
0·015
0·40
4·2
32
12
(La/Nd)CH
1·40
2·10
(La/Sr)PM
0·71
3·65
(La/Pb)PM
0·51
0·78
∗Calculated by the method of Bindeman et al. (1998); XAn = 0·57; T = 1173 K.
†Green (1994).
‡Sisson et al. (1994), Kd for Pb is calculated by interpolation.
§ Average value of the lower crust by Rudnick & Fountain (1995).
On the basis of petrological observations, quartz phenocrysts generally occur in felsic volcanic rocks without
hydrous phenocrysts in the volcanic front of the NE
Japan arc (Sakuyama, 1979). However, quartz does not
appear as a phenocryst in Akagi samples, even in the felsic
samples. Sakuyama (1979, 1983a, 1983b) determined the
H2O content of magmas based on the crystallization
sequence of phenocrysts. Following his definition, the
Akagi volcanic rocks belong to Type II–III, which allows
magmas to contain >3–4 wt % of H2O at >0·5 GPa.
This suggests that the H2O content of the Akagi magmas
is considerably higher than that of the Type I volcanoes,
typical of the volcanic front in NE Japan, which contain
<3 wt % of H2O.
Yamaguchi (1990) inferred that the groundmass in the
Akagi volcanic rocks has a significantly higher Al2O3
content than that of the adjacent Hotaka volcano. This
is because the plagioclase stability field becomes narrower
relative to the diopside stability field with increasing H2O
content in the magma. This observation is also consistent
with the hypothesis that the primary magma of Akagi is
enriched in H2O relative to other volcanoes in the
volcanic front in NE Japan.
The peculiar petrological and geochemical characteristics of Akagi volcano discussed above might be
explained by the development of the unusual tectonic
setting beneath Akagi volcanic area, where the Philippine
Sea plate overlaps onto the Pacific plate. Such a tectonic
setting might lead to a significantly larger amount of
water in the magma source region than that expected
from subduction of a single Pacific plate.
EVOLUTION OF MAGMA IN THE Ol
STAGE
Recently, magma mixing at Akagi has been discussed,
on the basis of the major element composition of phenocrysts and their textures, in the pumice from the Y stage,
by Horio & Umino (1995) and Umino & Horio (1998).
Those workers have suggested that high-T, less fractionated magma was periodically injected into a low-T,
silicic, mushy magma in a shallow magma chamber.
However, previous studies have not addressed the origin
of these two different magmas because of the limited
geochemical information available, especially isotope
compositions, for Akagi volcano.
On the basis of the major element compositions of
pyroxenes, the temperatures of the magmas in the Ol
stage were determined by using the QUILF program of
Andersen et al. (1993). Average compositions of the rims
of clinopyroxene and orthopyroxene were used for the
temperature estimation. Ferric–ferrous ratios were calculated to satisfy the stoichiometry of the pyroxenes. The
samples investigated from the Ol stage can be divided
into high-T (1000–1200°C) and low-T (<1000°C) groups,
and tend to have a negative correlation between calculated temperature and Sr isotopic composition (Table
8 and Fig. 12). This indicates that such a variation in
magma temperatures is not a consequence of the simple
cooling process in the shallow magma chamber.
In the Ol samples, melts are isotopically more depleted
than the coexisting phenocrysts (Fig. 9) as discussed
in the previous section. Such isotopic disequilibrium
observed could be explained only by a mechanical mixing
2322
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Fig. 11. Trajectories of AFC model calculation for Akagi volcanic rocks. The parameter r = Ma(t)/Mc(t) is defined as the assimilation rate, and
the numbers on the trajectories show the proportion of remaining magma (Mm/Mm0) defined by DePaolo (1981). The variation of trace element
ratios of Adatara volcano are calculated from Fujinawa (1992), and the Nd and Sr isotopic compositions are based on unpublished data. The
other parameters are compiled in Table 7. In the case where the assimilation rate is small (r = 0·2), the concentration ratios increase rapidly
without large isotopic differences. When the assimilation and fractionation are nearly balanced (r = 0·9), the trajectories appear to explain
moderately well the variation observed in the Akagi volcanic rocks.
between the isotopically enriched magma equilibrated
with the phenocrysts and the relatively depleted aphyric
magma. Mass-balance calculations, based on the trace
element composition and modal abundance of the phenocrysts, indicate that the Sr/Nd ratio in the glass of the
phenocryst-rich rock is lower than that in the aphyricprimary magma as a result of the equilibration of phenocrysts, especially plagioclase. If the aphyric-primary
magma has a more depleted isotopic signature than the
phenocryst-rich magma, the mixing curve forms a convex
hyperbola in the Sr and Nd isotope diagram (Fig. 13a).
That is, the Sr isotopic compositions of glasses are characteristically more depleted than those of the coexisting
phenocrysts. However, the difference in Nd isotopic
composition between them is smaller than the difference
in Sr isotopic composition. These observations, together
with model mixing, may suggest that the involvement of
depleted aphyric magma mixed with phenocryst-rich
magma was relatively small, and that isotopic re-equilibration among the phenocrysts and the mixed melt has
not been achieved.
To examine the above argument further, a model
calculation was carried out using the samples with remarkable isotope disequilibrium in Sr between melt and
phenocrysts (AK1012 and AK1102) representing the
high-T and the low-T group, respectively. The results
are presented in Fig. 13b, based on the parameters given
in Table 9. In this model, the isotope and trace element
compositions of the primary aphyric magma are assumed
to be those of the primary basaltic magma (a depleted
end-member, Pm) in the volcanic front of NE Japan
(Shibata & Nakamura, 1997). The isotopic compositions
of two end-members (Fm1 and Fm2) are represented by
the values of the most enriched phenocrysts in AK1012
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JOURNAL OF PETROLOGY
VOLUME 42
Table 8: Magmatic temperatures
obtained for Ol stage samples at 1 atm
and 2 kbar
Sample no.
1 atm
2 kbar
T (°C)
T (°C)
AK1108
808∗
814∗
AK1102
931±21
938±23
AK0910
1196±11
1206±10
AK1012
1069±20
1076±22
AK1201
1038±19
1045±21
AK-A
1068±12
1076±14
AK1010
1005±12
1012±14
∗Error not determined.
and AK1102, assuming that the phenocrysts in the enriched end-members were isotopically equilibrated with
their host melts, when the phenocrysts crystallized. The
Sr and Nd abundances in Fm1 and Fm2 were obtained
by the subtraction of phenocryst compositions from the
whole rock of the most mafic sample AK0910 in the Ol
stage. In this model, clinopyroxene and opaque phenocrysts are not considered. Because the Kd values of Sr
and Nd for clinopyroxene are almost identical (Green,
1994), and because of the extremely low concentrations
NUMBER 12
DECEMBER 2001
of these elements in the opaque minerals, very little
change in Sr/Nd is produced. The amounts of phenocrysts subtracted are assumed to be 30 wt % of plagioclase
for Fm1, and 40 wt % of plagioclase and 10 wt % of
hornblende for Fm2. After subtraction of the phenocrysts,
the resulting major element compositions are similar to
the whole-rock composition of the examined samples,
AK1012 and AK1102, except for Mg and Fe (e.g. calculated CSiO2Fm1 = 55·8, CSiO2Fm2 = 59·2 wt %, and the
whole rocks 56·4 and 59·5 wt %, respectively). In Fig.
13b, the calculated isotopic compositions of melts in
AK1012 and AK1102 are plotted near the fractions of
0·25 (Mm1) and 0·12 (Mm2) of the Pm component on
the mixing curves 1 and 2, respectively. The Sr isotopic
compositions show significantly more depleted signatures
than those of the whole rocks with less variation in
Nd isotopic composition. This is consistent with the
observation that the Sr isotopic compositions of most of
the glasses are characteristically more depleted than those
of the whole rocks and their constituent phenocryst
minerals. As shown in Fig. 9, the differences in both
Sr and Nd isotopic compositions between glasses and
phenocrysts are variable from sample to sample. This
may be attributed to the extent of mixing of the depleted
aphyric magma with the enriched porphyritic magma
and/or to the residence time after mixing in the magma
chamber, or to the evolution of the primary magma by
reaction with lower-crustal materials.
Although reverse zoning of pyroxenes is believed to
be evidence of magma mixing (i.e. Sakuyama 1979,
Fig. 12. Magmatic temperature vs Sr isotopic composition of whole rock, clinopyroxene and orthopyroxene phenocrysts in the Ol stage. The
magmatic temperature is determined using the QUILF program (Andersen et al., 1993).
2324
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Fig. 13. (a) Schematic illustration, in a Sr–Nd isotope diagram, of two-component mixing (after DePaolo & Wasserburg, 1979). K, Sr/Nd ratio
in component A relative to component B; F, fraction of component A in product magma. (b) Model calculation of mixing between the aphyric,
isotopically depleted melt (Pm) and the isotopically enriched melts (Fm1, Fm2). The chemical parameters are summarized in Table 9. The
numbers beside each curve represent the weight fraction of Pm in the mixed melt. The mixing trajectories are concave hyperbolae, and the K
values of Curve 1 and Curve 2 are 3·8 and 3·4, respectively. A relatively small amount of injection of the isotopically depleted primitive magma
(Pm) can produce the depleted melt part for Sr isotopes (Mm1 and Mm2) as a result of the mixing.
1981), such a signature was not observed in the sample
belonging to the high-T group, AK1012. If the depleted
primary magma (Pm), derived from a strongly metasomatized wedge mantle as a result of the double subduction, is less affected by assimilation processes in the
lower crust, the mode of phenocrysts could be smaller
than for anhydrous magma under the same temperature
conditions. This is because the liquidus temperature falls
with increasing H2O contents in melt (e.g. >1100°C at
>3 wt % H2O; Baker & Eggler, 1983). It may, therefore,
be reasonable that a depleted primary magma enriched
in water can cool without crystallization to a temperature
2325
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
DECEMBER 2001
Table 9: The compositions assumed for the end-member components in the mixing calculation
End-member
Pm
Fm1
Fm2
Mm1
Mm2
Sr (ppm)
290
230
160
245
173
87
Sr/86Sr
0·70430
Nd (ppm)
4·6
143
0·51282
Nd/144Nd
Nd
Sr/Nd
3·6
63
0·7072
0·70857
13·9
8·5
0·51236
0·51219
0·7063
11·6
0·5124
0·7077
8·0
0·5122
−5·5
−8·7
−4·5
−7·9
16·5
18·8
21·2
21·9
similar to that of the isotopically enriched end-member
(Fm1) of the high-T group. This consequently reduces
the temperature difference between the aphyric-depleted
and the enriched magma, making it difficult to form
reverse zoning in pyroxenes. On the other hand, the
low-T group represented by AK1102 commonly contains
pyroxenes with reverse zoning and hornblendes surrounded by pyroxene reaction rims. Under low-pressure
conditions, the hornblende stability field is sensitive to
magmatic temperature, and it dissociates to form crystal
aggregates of pyroxene and melts above >900–1000°C
(Boettcher, 1977). Thus, the occurrence of both reaction
rims in hornblende and reverse zoning of pyroxenes also
indicates that the temperature of the enriched endmember (Fm2) was significantly lower than that for Pm
and Fm1. Furthermore, the temperature of Fm2 was
raised by mixing with Pm. It is, however, difficult to
produce Fm2 directly from Fm1 by simple cooling processes associated with fractional crystallization in a shallow magma chamber, because of the significant difference
in isotopic composition between them (Fig. 12). It is
more likely that the high-T and low-T magmas behaved
independently in the shallow region through the Ol stage.
The temperature and isotopic differences between the
high-T and low-T magmas could have depended on the
difference in the degree of assimilation in the lower crust,
before their migration into the shallow magma chamber.
On the other hand, it may be possible to produce a highT magma from a low-T magma, when a single magma
batch such as low-temperature Fm2 is repeatedly injected
by a ‘hotter’ primary magma (Pm) in the shallow magma
chamber, resulting in rising magmatic temperature and
increasing isotopic depletion. However, a relatively large
discontinuity in isotopic composition and magmatic temperature exists between the high-T and low-T groups
(Fig. 12), indicating that two enriched magma reservoirs
have been involved in the formation of the Ol stage
magmas. The relatively small variations of isotopic compositions and temperature in each group may have been
caused by more complicated processes such as repeated
injections of relatively small amounts of new hot magma
into the enriched reservoirs in the shallow magma chamber.
It may be concluded that the isotopic variations and
disequilibrium between glass and coexisting phenocrysts
in the Ol stage were caused by two-component mixing
of relatively depleted aphyric magma and enriched
phenocryst-rich magma in the shallow magma chamber.
These two magmas evolved independently in the lower
crust before their mixing. This model is, therefore, different from that for the generation of calc-alkaline series
magmas proposed by Sakuyama (1979, 1981), in which
the original and its fractionated daughter magmas are
mixed, so that isotopic variations are not expected. Our
model is rather similar to the models of Eichelberger
(1978) and Takahashi (1986) based on petrological arguments, proposing a mixing of magma derived from
the mantle and a partially fused lower crust.
TEMPORAL ISOTOPIC VARIATION
OF AKAGI VOLCANO
The isotope signatures changed temporally during the
activity of Akagi volcano. The early stage volcanic rocks
are isotopically most depleted and the final stage becomes
depleted again, whereas the middle stage has more enriched signatures (Fig. 14). These temporal variations
are probably attributable to the evolution of the interface
(wall rock) between lower crust and the magma derived
from the mantle wedge beneath Akagi. When magma
starts to intrude into the lower crust, the margins of the
magma conduit could be chilled as a result of the significant temperature difference between the basaltic
magma and the lower crust. This could have led to less
interaction between the lower crust and the magma or
less assimilation by the magma resulting in a relatively
depleted isotopic signature at the early stage of Akagi
volcano. Repeated injection of magma should raise the
temperature of the wall rock near the magma conduit
and exceed the solidus temperatures of the surrounding
wall-rock materials in the lower crust. Thus, the volcanic
2326
KOBAYASHI AND NAKAMURA
EVOLUTION OF AKAGI VOLCANO
Fig. 14. Temporal variation of Sr, Nd and Pb isotopic compositions in the Akagi volcano.
rocks of the middle stage could have acquired the relatively enriched isotopic signature by lower-crustal assimilation relative to magmas of the early stage (Fig. 14).
In addition to the thermal evolution of the wall rock in
the lower crust, the interface between the magma and
the lower crust may have compositionally evolved, as a
result of the solidification of magma and the formation of
amphibole by the reaction between a water-rich primary
magma and pyroxene in the lower crust (see discussion
in previous section). Such an evolved interface could act
as an insulator that would prevent interaction between
the lower crust and the magma. Therefore, the isotope
signature in the final stage would become similar to that
of the early stage, with depleted isotope characteristics
relative to the middle stages (Fig. 14).
CONCLUSIONS
The evolution of magma to form Akagi volcano is schematically summarized in Fig. 15.
The major element compositions and the phenocryst
assemblages of Akagi volcano indicate that water content
in the magma is significantly higher than that of other
volcanoes in the volcanic front in the NE Japan arc. This
may be attributed to the unique geometry of subduction
beneath Akagi volcano. The Philippine Sea plate subducts
into the mantle wedge above the downgoing Pacific plate
underneath Akagi. Such double subduction of the oceanic
plates should provide a larger amount of water as a result
of the dehydration of hydrous minerals in the slabs, to
be contributed to the source region of magmas for Akagi
volcano, than in the case of single subduction of the
Pacific plate, which produced other volcanoes in the
volcanic front in NE Japan.
On the other hand, multi-isotope and trace element
systematics of Akagi volcano clearly require interaction
between the water-rich magma and the lower crust along
with amphibole fractionation. Such assimilation–
fractional crystallization in the lower crust may largely
have been induced by a fall in the solidus temperature
of the lower crust as a result of the higher water content
in the primary magma of this volcano relative to that of
other volcanoes in the volcanic front. The AFC model
calculation suggests that the latent heat of crystallization
and the heat of fusion for the wall rock were nearly
balanced in the lower crust, implying that assimilation
proceeded easily because of the high temperature in the
lower crust and the high water content in the magma.
Isotope disequilibrium was observed between phenocrysts formed in the shallow magma chamber and the
coexisting melts. Furthermore, the apparent magma temperatures correlate with Sr isotope compositions but to
a lesser extent with Nd isotope compositions. These
observations suggest that at least two recognizable and
distinct magmas, which had been separately evolving in
the lower crust, existed in the shallow magma chamber
in a single stage. Then, the magmas were mixed independently with injections of isotopically depleted aphyric magmas. Although the other stages were not examined
for phenocryst geochemistry in this study, it may be also
possible, on the contrary, that the depleted magma is
mixed by the injection of enriched magma. Such a
difference should be attributable to the residence time of
magma, and the evolution of the interface between the
magma and the wall rock in the lower crust. Therefore,
we conclude that the evolution of Akagi volcano has
been strongly controlled by the assimilation process in
the lower crust induced by the water-rich magma compositions.
2327
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 12
DECEMBER 2001
Fig. 15. Schematic representation of the magma differentiation process, including assimilation and amphibole fractionation by water-rich magma
in the lower crust under Akagi volcano. Overlapping of the Pacific and Philippine Sea plates may be an important mechanism for supplying
fluid to form the H2O-rich primary magmas. Finally, fractional crystallization and repeated magma mixing processes in a shallow magma
chamber produced the various geochemical characteristics of the Akagi volcanic rocks.
As mentioned repeatedly in this paper, a primary factor
in enhancing assimilation in the lower crust is the higher
water content in the primary magma than that in magmas
producing the surrounding volcanoes in the volcanic
front. If this is the case, the along-arc Sr isotope variation,
which decreases on either side of the ‘unique’ Akagi
volcano in the volcanic front, may indicate a decrease
in the extent of assimilation in the lower crust as a result
of a decrease in the water content of the primary magmas.
This may be due to the location of Akagi volcano, which
stands above the top of a ‘tongue’ of the Philippine Sea
plate. The position of this feature is shown by the
isodepths to the subducting plate in Fig. 1. This would
further increase the water content in the source region
of the primary magma by the dehydration of the Philippine Sea plate, in addition to the water from the Pacific
plate. If this is true, it may be possible to more clearly
deduce the extent of the Philippine Sea plate in the
mantle wedge of the Pacific plate and the role of water
in the generation of island-arc magmas by extending
detailed trace element and multi-isotope systematics to
the volcanoes in central Japan as has been performed
for Akagi volcano in this study.
ACKNOWLEDGEMENTS
We thank T. Shibata, M. Yoshikawa and A. Makishima
for their technical help in isotopic analyses and ICP-MS
analysis. Thanks are also due to our many colleagues
with whom we have discussed various issues during the
course of this study. We are deeply indebted to Professor
I. Kushiro for his encouragement. Professor I. Moriya is
also sincerely thanked for geological information on the
study area. Dr G. E. Bebout is also acknowledged for
improving the manuscript. This research is supported by
JSPS Fellowships for Japanese Junior Scientists (K.K.),
and by the Ministry of Education, Science, Sports and
Culture of the Japanese Government (Monbukagakusho)
and the Japanese Society for the Promotion of Science
( JSPS) (E.N.).
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