JOURNAL OF PETROLOGY
VOLUME 55
NUMBER 6
PAGES 1083^1128
2014
doi:10.1093/petrology/egu020
Progressive Interaction between Dry and Wet
Mantle during High-temperature Diapiric
Upwelling: Constraints from Cenozoic
Kita-Matsuura Intraplate Basalt Province,
Northwestern Kyushu, Japan
TETSUYA SAKUYAMA1*, SHUN’ICHI NAKAI2,
MASAKO YOSHIKAWA3, TOMOYUKI SHIBATA3 AND
KAZUHITO OZAWA4
1
INSTITUTE FOR RESEARCH ON EARTH EVOLUTION, JAPAN AGENCY FOR MARINE^EARTH SCIENCE AND
TECHNOLOGY, YOKOSUKA 237-0061, JAPAN
2
EARTHQUAKE RESEARCH INSTITUTE, UNIVERSITY OF TOKYO, TOKYO 113-0032, JAPAN
3
INSTITUTE FOR GEOTHERMAL SCIENCES, UNIVERSITY OF KYOTO, NOGUCHIHARA, BEPPU CITY, OITA
PREFECTURE 874-0903, JAPAN
4
DEPARTMENT OF EARTH AND PLANETARY SCIENCE, GRADUATE SCHOOL OF SCIENCE, UNIVERSITY OF TOKYO,
TOKYO 113-0033, JAPAN
RECEIVED DECEMBER 25, 2013; ACCEPTED MARCH 31, 2014
Intra-plate Cenozoic volcanism in Kita-Matsuura, northwestern
Kyushu, Japan, shows systematic spatio-temporal changes in geochemistry that can be explained by partial melting followed by melt
segregation in a region of upwelling mantle. We have examined the
thermal and melting history of the upwelling mantle by quantitatively estimating melt water contents and melting conditions. The
water content of a spectrum of primary melts is estimated to range
from 0·5 to 1·5 wt % based on a combination of a plagioclaseliquid and olivine-saturated liquid geohygrometers and MELTS calculations. The estimated melt segregation temperature ranges from
1330 to 15008C, at pressures from 1·7 to 2·8 GPa under hydrous conditions. Melting temperature and pressure decreased with time,
whereas the water content of the primary melts increased.
Corresponding temporal decreases in high field strength element
(HFSE) abundances and HFSE/large ion lithophile element
(LILE) ratios require progressive melt extraction and aggregation
KEY WORDS: Kita-Matsuura; Japan; intraplate back-arc volcanism;
open-system melting
*Corresponding author. Telephone: þ81-46-867-9785.
þ81-46-86-9625. E-mail: [email protected]
ß The Author 2014. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oup.com
Fax:
from a melting mantle with a continuous and gradually increasing
input of H2O-rich fluid or melt into the melting system. The estimated isotope composition of influxed fluid lies on a mixing line between the sediment and altered oceanic crust of the Philippine Sea
Plate, with strong affinity to the sediment composition. Based on
the temporal variation of the magmas and the melting model, we propose small-scale upwelling (c. 70 km in diameter) of a dry mantle
peridotite that interacts progressively with the overlying wet mantle
wedge. The wet mantle wedge was previously hydrated by fluids
from sediments from the subducted Philippine Sea Plate, whereas
the deep and dry mantle could have been derived from the mantle beneath the subducted Pacific Plate through a slab window.
JOURNAL OF PETROLOGY
VOLUME 55
I N T RO D U C T I O N
Cenozoic back-arc volcanism has occurred extensively, but
sporadically, along 1000 km of the eastern margin of the
Eurasian Plate. The back-arc region, where the subducting
slab surface is deeper than 300 km, is characterized by
high surface heat flow (Currie & Hyndman, 2006), upper
mantle with slow seismic velocity (e.g. Fukao et al., 1992)
and high electrical conductivity (Ichiki et al., 2006).
Various mantle upwelling models have been postulated to
explain the association of these geophysical characteristics
with the volcanic activity. These include the following: (1)
hot mantle plumes originating from the upper mantle^
lower mantle or core^mantle boundary (Nakamura et al.,
1990; Zeng et al., 2011); (2) upwelling of hydrous hot
mantle plumes originating from the 410 km discontinuity
(Zhao et al., 2007; Kuritani et al., 2009); (3) wet and cold
plume upwelling (Richard & Iwamori, 2010); (4) a
moving hot region distinct from a hotspot (Miyashiro,
1986); (5) asthenospheric upwelling induced by thickening
of a stagnant slab (Zou et al., 2008); (6) lateral mantle extrusion induced by the Indo-Asian continental collision
(Liu et al., 2004). A role for mantle upwelling in causing
back-arc volcanism, as invoked in these models for eastern
Asia, has also been suggested for other areas where backarc volcanism occurs, such as New Zealand (Rafferty &
Heming, 1979) and Patagonia (Ignacio et al., 2001). To identify the actual upwelling process among the diverse
models that have been proposed, the spatial and temporal
relationships between the various processes causing volcanism in each back-arc region need to be quantified.
Back-arc volcanoes that occur on continents, especially if
they are extinct and young, have an advantage over oceanic back-arc volcanoes, as their entire volcanic sequence
can be exposed in gullies formed by erosion of the flanks,
and thus can be much more easily sampled at a high temporal resolution. Moreover, continental back-arc volcanoes
generally show primitive petrological characteristics, and
although the volcanic activity is often spread over a wide
area at different eruptive centres, each volcano has a limited extent. As a result, petrological and geochemical data
for volcanic rocks in this tectonic setting that are well constrained spatially and temporally can provide spatial and
temporal information on mantle upwelling and melting,
as long as the effects of crustal processes can be removed.
However, only a few studies have exploited this advantage
of continental back-arc volcanism (e.g. Iwamori, 1991;
Sakuyama et al., 2009).
Sakuyama et al. (2009) made systematic geological, geochemical, and chronological investigations of Cenozoic
intraplate volcanism in the Kita-Matsuura area, southwestern Japan, which they interpreted in terms of mantle upwelling beneath this region. Their model was, however,
far from quantitative as they assumed anhydrous melting
without hydrous material input, which caused difficulties
NUMBER 6
JUNE 2014
in the estimation of mantle potential temperature and the
explanation of the observed fractionation between fluidmobile and fluid-immobile elements. In this study, new
mineral chemical data for plagioclase phenocrysts and
spinel in the basalts, whole-rock trace element data measured by inductively coupled plasma mass spectrometry
(ICP-MS), and whole-rock Sr^Nd^Pb isotope data measured by thermal ionization mass spectrometry (TIMS)
are presented. These data are combined with the previous
results of Sakuyama et al. (2009) to estimate quantitatively
the temporal and spatial changes in melting conditions,
including water content, pressure, and temperature. The
results successfully place stricter constraints on the melting
process and allow us to accurately constrain the mantle potential temperature, which has important geodynamic implications for upwelling in the back-arc region of Kyushu,
southwestern Japan.
G EOLO GY A N D P R EV I O U S
ST U D I E S OF T H E
K I TA- M AT S U U R A B A S A LT S
Northwestern Kyushu is located between two Cenozoic
marginal basins: the Japan Sea and Okinawa Trough. The
Pacific Plate is estimated to have been subducting beneath
northeastern Asia since 60 Ma (Mu«ller et al., 2008), followed by initiation of subduction of the Philippine Sea
Plate at the time of initiation of Japan Sea opening at 25
Ma (Kimura et al., 2005). Since the opening of the Japan
Sea between 25 and 15 Ma (Tamaki et al., 1992), the
Philippine Sea Plate has been subducting northwestward
under southwestern Japan. North of this area the Pacific
Plate is subducting westward, and the top of the slab is at
a depth of 400^600 km beneath southwestern Japan. The
subduction angle of the younger part of the Philippine
Sea Plate is shallow (128) beneath the Chugoku area,
whereas that of the older part of the Philippine Sea Plate
is steep (608) beneath Kyushu. Opening of the Okinawa
Trough initiated in Late Miocene times and its northern
extension continues to the Beppu-Shimabara graben in
western Kyushu (Letouzey & Kimura, 1985), in which the
active Unzen volcano is located.
The most voluminous Cenozoic basaltic volcanism in
southwestern Japan occurred at the end of the Miocene in
the Kita-Matsuura area of northwestern Kyushu associated with a compressional stress field (Sakuyama, 2010).
The Kita-Matsuura area is located 150 km to the backarc side of the current volcanic front in Kyushu, which is
400 km from the trench. This basaltic volcanism began
after a dormant period of 5 Myr following basaltic and
andesitic volcanism on Hirado island at 15 Ma during
the Middle Miocene when the Japan Sea stopped opening
(Jolivet et al., 1994).
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MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Sakuyama et al. (2009) investigated four volcanic sections that are inferred to have originated from single
volcanic centers and together encompass the entire KitaMatsuura basalt succession (from west to east the Hirado,
Senryu, Ishimori, and Kunimi sections; Fig. 1). Volcanic
activity started at 9·5 Ma in the Hirado section and continued intermittently for 2 Myr in all of the sections
(Fig. 2a). In addition to the Kita-Matsuura area,
Cenozoic basaltic volcanism also occurred on several of
the islands to the west and north of the Hirado section:
Takushima, Ikitsuki, and Azuchi-Oshima (Fig. 1a). The
Takushima and Ikitsuki basalts range in age from 9·5 to
8·7 Ma and from 8·5 to 6·5 Ma, respectively, and are similar to the basalts in the Kunimi section in terms of their
high contents of Na2O þ K2O and low MgO at
SiO2 ¼50 wt % (Uto et al., 2004), which suggest that they
are cogenetic with the Kita-Matsuura basalts. Below, we
summarize the results of Sakuyama et al. (2009) upon
which this study builds.
Spatial and temporal variations in the
Kita-Matsuura basalts
After evaluating the effects of crustal processes (magma
mixing, crustal assimilation, crystal fractionation),
Sakuyama et al. (2009) showed that the diversity of basalt
chemistry can be represented by three liquid lines of descent with distinct trends on oxide^oxide variation diagrams: low- (47^50 wt %), medium- (49^52 wt %), and
high-SiO2 (51^54 wt %) groups. The groups can also be
defined based on their incompatible element ratios,
whereby Nb/Y, Zr/Y, and Nb/Th decrease from the lowto high-SiO2 groups (Fig. 2b and c). Most of the samples
in the Hirado and Senryu section are classified as belonging to the medium- and high-SiO2 groups (Fig. 2a),
whereas low-SiO2 group lava is more common in the eastern sections, such as at Ishimori and Kunimi. In the
Hirado, Senryu, and Ishimori sections, lavas sequentially
change their composition to the higher-SiO2 group
(Figs 1b^d and 2a); in particular, from the medium- to
high-SiO2 group in Hirado, from the low-, through
medium-, to high-SiO2 group in Senryu, and from the
low- to high-SiO2 group in Ishimori.
Mantle upwelling model and problems that
need to be solved
Fig. 1. (a) Simplified map around Kyushu, southwestern Japan. Bold
continuous curves represent plate boundaries. Okinawa Trough and
Beppu-Shimabara Graben are indicated by bold dotted curves. (b)
Distribution of the Kita-Matsuura basalts in northern Kyushu. (c)
Geological cross-sections through the Kita-Matsuura basalts along
the lines A^B, C^D, E^F, and F^G shown in (b). The vertical scale
is twice the horizontal scale. The four cross-sections pass through
the four areas studied: Hirado, Senryu, Ishimori, and Kunimi on
the A^B, C^D, E^F, and F^G sections, respectively. These are
used to clarify the simplified sequence of magmatic groups
Sakuyama et al. (2009) found that the melt segregation
pressure decreases from 3 GPa for the low- to 1·5 GPa
Fig. 1 Continued
after Sakuyama et al. (2009). Low^medium- and medium^high-SiO2
groups, which are samples classified into both a low- and mediumSiO2 group and a medium- and high-SiO2 group, respectively, are
omitted from this figure for simplicity [see Sakuyama et al. (2009) for
the detailed magmatic sequence]. masl, meters above sea level; filled
triangle represents 0 masl.
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JOURNAL OF PETROLOGY
VOLUME 55
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for the high-SiO2 group under anhydrous conditions.
Following Sobolev & Shimizu (1993) decreases of Nb/Y
and Zr/Y from the low-, through medium-, to high-SiO2
groups (Fig. 2b and c) were attributed to an increase of
melting degree by progressive near-fractional melting of a
primitive or enriched mantle source by applying a melting
model open to only output; the temporal decrease in melting pressure and increase of melting degree in every section originated in a mantle upwelling with progressive
melting and melt segregation.
However, several critical issues were not fully addressed
by Sakuyama et al. (2009), including the following: (1) the
effect of water in mantle melting was not examined; (2)
only the ranges of primary melts were estimated because
of uncertainty in the source mantle composition in
terms of olivine forsterite content [Fo# ¼ 100 Mg/
(Mg þ Fe2þ)]; (3) melting pressure was only approximately
constrained because of the uncertainty in the primary
melt compositions; (4) melting temperature (or mantle
potential temperature) was not estimated; (5) only fluidimmobile high field strength elements (HFSE) were used
in trace element modelling and fluid-mobile elements
were not considered. To resolve issues (1), (2), (3), and (4),
we have obtained new mineral chemical data; to address
issue (5), we obtained a highly accurate trace element and
isotope dataset by ICP-MS and TIMS.
A N A LY T I C A L M E T H O D S
We selected 15 samples from the four stratigraphic sections
through the Kita-Matsuura basalts for trace element and
isotope analysis: three samples were selected from each of
the Kunimi, Ishimori, and Hirado sections and six samples were selected from the Senryu section. Samples were
selected to include at least the initial, middle, and terminal
basaltic activities in each section. Olivine phenocrysts or
microphenocrysts in all the samples selected for analysis
were fresh without iddingsite in the core. Totals of 10
major elements determined by X-ray fluorescence (XRF)
for all of the samples were 498 wt %, which is suggestive
of a minimum effect of subaerial modification on bulkrock chemical compositions.
Whole-rock major element contents for most samples in
this study were taken from Sakuyama et al. (2009), except
for samples 03101930, 03102481, 031027c9, and 031026a8,
which were newly analysed by XRF using a Philips
PW-1480 system at the Department of Earth and Planetary
Fig. 2. Summary of the temporal and spatial variations in the KitaMatsuura basalts based on Sakuyama et al. (2009) and Uto et al. (2004).
Fig. 2 Continued
(a) Temporal changes of activity in the studied sections according to
Sakuyama et al. (2009). Data for Ikitsuki and Takushima are from
Uto et al. (2004). Range of eruption ages are indicated by filled bars.
(b) Relationship between estimated ranges of segregation pressure
and Nb/Y ratio for each group of basalts. (c) Zr/Y vs Nb/Y for each
basalt group after Sakuyama et al. (2009).
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SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Science, University of Tokyo. Rock samples were sliced into
chips using a diamond cutter, polished with a diamond,
and powdered in an agate mortar. Detailed preparation
and analytical methods for major element composition
have been described by Sakuyama et al. (2009).
Twenty-eight trace elements [Rb, Ba, Th, U, Pb, Ta, Nb,
Sr, Zr, Hf, Y, and the rare earth elements (REE)] were
determined for 15 samples of the Kita-Matsuura basalt by
ICP-MS at the Earthquake Research Institute, the
University of Tokyo. Selected rock powders were weighed
into Savillex Teflon vials with hydrofluoric acid and perchloric acid. The vials were heated to about 1808C for 2
days. The decomposed samples were then evaporated. The
dried samples were dissolved in 20% nitric acid to ensure
that there was no precipitation. A part of the solution was
diluted in 1·5% nitric acid such that the final dilution
factor was 6000. Indium and bismuth were added as internal standards to correct for machine drift and matrix effects. The reproducibility of the analyses for all the
elements was better than 5% repeated standard deviation
(RSD). Our abundance results on the GSJ standard basalt
JB-1a agreed with those of the reference values to within
10%. The results are listed in Table 1. We checked the consistency between the ICP-MS analyses and those determined by XRF by Sakuyama et al. (2009) (see Appendix).
Mineral compositions of plagioclase, spinel, and olivine
were analysed by electron microprobe (JEOL JCMA733MKII) at the University of Tokyo. The analytical procedures have been described by Nakamura & Kushiro (1970),
and used the Bence & Albee (1968) correction method.
Accelerating voltage, beam current, and counting time were
15 kV, 12 nA, and 30 s, respectively. The Fo# and NiO content of olivine were determined at 25 kV with a 20 nA beam
current and 30 s counting time using a ZAF correction.
Rock samples for Sr^Nd^Pb isotope analysis were sliced
into chips using a diamond saw, polished with abrasive
powder to remove the tracks of the cutter, and washed in
an ultrasonic bath. After the chips were crushed to a few
millimeter-sized fragments in a tungsten carbide mortar,
the fragments were ground in an automatic agate mortar
for 2 h. The Sr, Nd, and Pb isotopic compositions of the samples were determined by TIMS at the Institute for
Geothermal Sciences, Kyoto University. Details of the analytical procedure for chemical separation and mass spectrometry for Sr, Nd, and Pb isotope determination have
been given byYoshikawa et al. (2001) and Shibata et al. (2003).
W H O L E - RO C K T R AC E E L E M E N T
A N D I S OTOP E A N D M I N E R A L
C H E M I C A L DATA
Whole-rock trace element composition
Figure 3a^d shows trace element patterns normalized to
normal mid-ocean ridge basalt (N-MORB) (Pearce &
Parkinson, 1993) for the four sections of the KitaMatsuura basalt that were sampled. The three basalt
groups (high-, medium-, and low-SiO2 groups) are labeled
in each section.
The REE patterns are smooth and show enrichment in
light REE (LREE) with significant variations in the overall slope and the LREE abundances. There is a smooth decrease from Ba to U, and the relative normalized
abundances of Ba, Th, and U are similar (Fig. 3a^d). No
basalt sample has an Eu anomaly. Positive Ba, Pb, and Ti
anomalies and weak positive Sr anomalies are generally
observed. As the concentrations of Nb decrease, the Pb enrichment relative to Ce or Pr and the Nb depletion relative
to Th or K increase (Fig. 3e). Generally, the abundances
of moderately to highly incompatible trace elements in the
eastern sections (Kunimi and Ishimori) are higher than
those in the western sections (Senryu and Hirado).
The abundance of incompatible trace elements and Nb/
Th and Nb/Y decrease with decreasing age in each section
(Sakuyama et al., 2009), although they show relatively
wide scatter over a small timescale (less than a few million
years). Both the REE abundance and slope of REE normalized to N-MORB decrease with time (Fig. 3a^d). The
latter is clearly shown by the temporal variations of Ce/
Yb and Sm/Yb, which decrease with time in each section
(Fig. 4c and d).
The abundance of incompatible trace elements, and the
middle REE (MREE)/heavy REE(HREE) and the
LREE/HREE ratios are intimately related to the SiO2
grouping. Weak negative Nb and Ta anomalies occur in
all the high-SiO2 group basalts and in some of the
medium-SiO2 group basalts, but not in the low-SiO2
group basalts (Fig. 3a^d). The abundance of incompatible
trace elements and MREE/HREE, LREE/HREE, Zr/Y,
Nb/Th, and Nb/Ce decrease from the low- to high-SiO2
groups (Figs 3 and 4).
Chemical compositions of spinel inclusions
in olivine phenocrysts and plagioclase
phenocrysts or microphenocrysts
Olivine phenocrysts in less fractionated basalts (FeO*/
MgO 51·5) contain spinel inclusions 10^30 mm in diameter, although spinel microphenocrysts are rare.
Representative major element compositions of spinel inclusions in the least differentiated samples for each SiO2
group are given in Table 2. The spinel Cr# [100 Cr/
(Cr þAl)] ranges between 40 and 60 (Fig. 5a) and the
YFe [100 Fe3þ/(Cr þAl þ Fe3þ)] averages 15, except for
spinel in Fe-rich olivine (Fig. 5b). As the Fo# of the host
olivine decreases, the Cr# of the spinel inclusions is
almost constant for the low-SiO2 group and decreases
slightly for the high-SiO2 group (Fig. 5a). This relationship
between the Cr# of spinel inclusions and the Fo# of host
olivine is commonly observed in spinel inclusions in olivine
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Table 1: Major and trace element concentrations for representative Kita-Matsuura basalts, standards and estimated primary
melts
Section:
Hirado
Hirado
Hirado
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Ishimori
Ishimori
Ishimori
Basalt group:
high
high
medium
high
high
high
medium
medium
low
high
high
low
low
8·1
8·84
6·37
6·48
6·54
6·56
6·59
7·68
5·42
7·33
7·74
7·11
K–Ar age1 (Ma): 7·29
Name:
Kunimi
03101930 03101924 03101817 02110723 02110934 02111276 02111173b 03102155 03102145 031029if7 03102481 03102475 031027e4
Major elements (wt %)
SiO2
50·86
52·80
50·07
50·83
49·31
50·75
47·47
49·45
46·44
51·41
51·16
48·83
TiO2
1·56
1·32
1·55
1·16
1·20
1·54
1·46
1·91
2·17
1·20
1·47
2·79
1·82
Al2O3
14·85
15·32
17·97
14·85
14·08
16·48
13·96
16·13
17·13
15·73
15·43
16·74
13·84
Fe2O3*
11·3
11·34
46·46
12·38
10·45
9·13
11·73
10·54
10·08
10·94
9·87
8·82
10·39
12·46
MnO
0·17
0·15
0·16
0·16
0·17
0·16
0·17
0·16
0·16
0·16
0·15
0·14
0·17
MgO
7·38
7·08
6·59
9·0
10·42
6·65
9·62
7·08
7·64
7·63
7·67
4·48
10·81
CaO
9·12
8·23
8·64
8·80
9·10
9·77
10·91
9·79
10·09
9·77
8·36
7·42
11·48
Na2O
2·92
2·70
2·87
2·68
2·39
3·07
2·46
2·74
2·17
2·60
2·81
3·07
1·93
K2O
0·73
1·02
1·64
0·53
0·72
0·81
0·79
1·28
1·15
1·01
1·09
1·86
0·96
P2O5
0·22
0·24
0·41
0·15
0·21
0·31
0·27
0·41
0·41
0·25
0·36
0·51
0·27
Total
99·88
99·31
99·04
99·90
98·13
99·61
98·07
98·82
98·66
98·60
98·88
98·33
99·06
Trace elements (ppm)
Rb
Ba
16·6
145
26·0
226
42·2
471
12·7
132
15·4
233
11·4
335
16·1
373
27·3
404
23·6
1000
22·7
345
24·1
300
51·4
30·3
484
418
Th
2·19
3·33
6·03
1·75
2·55
3·09
3·92
4·05
3·60
3·35
3·52
6·44
5·46
U
0·41
0·72
0·87
0·38
0·56
0·63
0·81
0·91
0·71
0·69
0·76
1·44
1·07
Ta
0·46
0·81
1·12
0·34
0·58
0·97
0·77
1·51
1·62
0·65
1·12
2·65
Nb
7·73
Pb
Sr
Zr
Hf
Y
2·91
346
94·6
2·67
19·6
12·9
4·55
18·9
7·24
308
502
120
151
3·20
19·3
3·80
17·8
5·75
3·63
267
69·9
2·06
18·9
10·1
3·04
336
84·3
2·37
16·6
16·3
4·10
431
14·8
3·28
453
147·1
3·82
19·9
98·8
2·74
19·2
24·8
4·45
26·5
2·82
508
621
173
155
4·51
21·0
4·04
10·9
4·29
449
80·8
2·40
18·7
4·23
44·2
4·43
397
447
131
222
3·46
1·52
25·4
2·76
506
99·8
5·53
2·82
18·8
20·6
19·6
26·6
17·7
La
12·2
17·6
30·9
10·2
14·5
20·9
23·0
26·2
25·2
24·0
22·2
38·1
27·0
Ce
28·6
35·8
65·5
20·4
29·5
43·6
43·3
53·1
54·9
40·6
45·8
76·9
51·4
Pr
Nd
3·81
16·5
4·56
18·5
7·48
28·3
2·94
12·7
3·57
14·8
5·25
21·3
5·04
20·3
6·46
26·1
6·63
26·8
5·02
20·3
5·40
21·9
9·03
35·4
5·90
23·6
Sm
4·10
4·37
5·11
3·38
3·48
4·76
4·44
5·67
5·56
4·32
4·86
7·25
4·84
Eu
1·39
1·44
1·63
1·21
1·18
1·62
1·48
1·88
1·83
1·49
1·67
2·28
1·61
Gd
4·55
4·56
4·39
4·05
3·75
4·81
4·55
5·45
5·21
4·41
4·97
6·90
4·59
Tb
0·70
0·70
0·63
0·65
0·57
0·72
0·67
0·80
0·74
0·65
0·72
1·02
0·67
Dy
4·19
4·20
3·73
3·95
3·47
4·23
4·03
4·59
4·29
3·94
4·29
5·87
3·90
Ho
0·83
0·82
0·73
0·79
0·69
0·82
0·79
0·88
0·82
0·80
0·85
1·13
0·75
Er
2·28
2·22
2·03
2·24
1·88
2·28
2·15
2·39
2·14
2·21
2·28
2·99
2·00
Tm
0·31
0·31
0·29
0·31
0·26
0·32
0·30
0·33
0·29
0·30
0·32
0·41
0·27
Yb
1·99
2·00
1·91
2·04
1·69
2·06
1·94
2·08
1·85
1·93
2·05
2·57
1·68
Lu
0·28
0·29
0·28
0·29
0·24
0·30
0·27
0·29
0·26
0·26
0·29
0·29
0·36
(continued)
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Table 1: Continued
Section:
Kunimi
Kunimi
Basalt group:
low
low
K–Ar age1 (Ma):
7·26
7·76
Name:
031027c9
031026a8
Standard
Estimated primary melt
Fo#:
JB-1a
JB-1a
(this study)
(reference value)
low
medium
high
90
90
90
Major elements (wt %)
SiO2
49·20
47·93
45·89
47·02
TiO2
2·29
3·13
1·56
1·24
49·01
1·07
Al2O3
16·96
16·79
11·89
11·88
12·59
Fe2O3*
10·89
10·71
13·00
11·56
11·38
MnO
0·15
0·17
0·16
0·17
0·17
MgO
6·08
4·72
16·36
16·02
15·16
CaO
9·17
7·12
9·87
9·28
8·14
Na2O
2·73
3·10
1·66
2·10
2·14
0·65
K2O
1·32
2·03
0·82
0·67
P2O5
0·46
0·61
0·23
0·23
Total
99·07
98·61
100
100
0·18
100
Trace elements (ppm)
Rb
Ba
30·2
464
49·7
618
35·7
39·2
461
504
26·4
262
13·6
316
10·1
106
Th
5·99
7·29
9·53
9·03
3·37
3·32
1·40
U
1·23
1·44
1·81
1·57
0·75
0·68
0·31
Ta
1·73
2·88
1·63
1·93
1·37
0·65
0·28
25·3
26·9
6·70
6·76
Nb
Pb
27·7
3·85
48·5
6·03
22·8
2·34
Sr
603
520
423
442
288
Zr
153
246
128
144
118
3·52
3·41
Hf
3·90
6·10
3·05
12·5
2·78
384
83·9
2·32
4·6
2·91
220
56·1
1·66
Y
20·2
25·9
20·9
24·0
15·3
16·3
La
34·6
46·5
43·0
37·6
20·0
19·5
8·1
Ce
65·1
97·3
72·5
65·9
40·6
36·7
16·4
Pr
Nd
7·46
29·1
11·06
9·2
7·3
43·0
29·3
26·0
4·84
19·2
4·27
17·2
15·3
2·36
10·3
Sm
5·77
8·29
5·95
5·07
4·03
3·76
2·73
Eu
1·93
2·60
1·47
1·46
1·36
1·25
0·99
Gd
5·34
7·27
5·30
4·67
3·85
3·86
3·28
Tb
0·76
1·04
0·67
0·69
0·57
0·57
0·53
Dy
4·39
5·80
5·37
3·99
3·30
3·42
3·20
Ho
0·84
1·09
0·80
0·71
0·63
0·67
0·64
Er
2·22
2·81
2·90
2·18
1·66
1·82
1·81
Tm
0·31
0·38
0·31
0·33
0·23
0·26
0·25
Yb
1·99
2·42
2·70
2·10
1·43
1·65
1·65
Lu
0·28
0·34
0·29
0·33
0·20
0·23
0·23
*Total iron given as Fe2O3.
1
K–Ar ages are the values that were calculated by regressing separately for each section to estimate ages for lava flows
without age determination [see details given by Sakuyama et al. (2009)]. Major element compositions of all the samples in
this table were previously reported by Sakuyama et al. (2009) except for 03101930, 03102481, 031027c9, and 031026a8,
which were newly analysed for this study by the same analytical method as that of Sakuyama et al. (2009).
Fo#, Fo content of equilibrium mantle olivine.
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Representative major element compositions of plagioclase phenocrysts or microphenocrysts are given in
Table 3. Plagioclase phenocrysts in the low- and mediumSiO2 groups are euhedral to subhedral and 0·5^3 mm in
average size, reaching a maximum of 1cm. The anorthite
(An) [100 Ca/(Ca þ Na)] content in the cores varies
from 60 to 80, and in the rims from 40 to 70 (Fig. 6a and
b). The range of An contents in the groundmass plagioclase is systematically lower than that of the phenocrysts
(Fig. 6a and b).In relatively differentiated lava samples
(bulk FeO*/MgO41·25) normally zoned plagioclase
forms crystal aggregates with olivine and clinopyroxene.
The olivine and clinopyroxene in these aggregates have
cores with lower Fo# (570) and Mg# [100 Mg/
(Mg þ Fe2þ)] (575), respectively, than those of isolated
olivine and clinopyroxene phenocrysts.
Large plagioclase phenocrysts are rare in the high-SiO2
group (51% by volume), but euhedral to subhedral normally zoned microphenocrysts are common. These vary
from 100 to 200 mm in length and are more than twice the
average size of the groundmass plagioclase. The An contents of the microphenocrysts varies from 65 to 87, with a
clear mode at 65^70 (Fig. 6c). Some plagioclase microphenocrysts form crystal aggregates with olivine and clinopyroxene. The olivine and clinopyroxene in these aggregates
have cores with lower Fo# (570) and Mg# (580), respectively, than the cores of isolated olivine and clinopyroxene phenocrysts. The range of An contents in the
groundmass plagioclase is 60^70, which is slightly lower
than that of the microphenocrysts.
Whole-rock Sr^Nd^Pb isotope compositions
Fig. 3. (a^d) Trace element patterns normalized to N-MORB
(Pearce & Parkinson, 1993) for the four studied sections of the KitaMatsuura basalts. Areas encircled by broken lines represent the
entire range of Kita-Matsuura basalt samples analysed in this study.
In each panel, the three basalt groups (high-, medium-, and lowSiO2 groups) are distinguished. (e) Trace element patterns normalized to sample 031026a8 (low-SiO2 basalt) for high- (continuous
line), medium- (dashed line), and low-SiO2 (grey field) groups.
phenocrysts and can be attributed to fractional crystallization of olivine and spinel (Arai, 1994). The host olivines exhibit bell-shaped Fe^Mg zoning patterns, which indicate
that the Fo# around the spinel inclusions was not modified during cooling as the samples were obtained from
thin lava flows or the surface of thick lava flows.
The Sr^Nd^Pb isotopic compositions are reported in
Table 4 and plotted in Fig. 7. These are in the range of
0·70362^0·705479 in 87Sr/86Sr, 0·512584^0·512866 in
143
Nd/144Nd, 18·036^18·451 in 206Pb/204Pb, 15·493^15·715 in
207
Pb/204Pb, and 38·326^39·125 in 208Pb/204Pb. All the
samples from this study plot within the range of other
Cenozoic basalts from eastern Asia (Fig. 7). The isotopic
compositions of sediments on the Pacific Plate (PAC) and
the Philippine Sea Plate (PHS) are also plotted in Fig. 7
(Hickey-Vargas, 1991; Plank & Langmuir, 1998; Shimoda
et al., 1998; Hauff et al., 2003; Plank et al., 2007). They plot
in the more enriched Sr^Nd isotope quadrant and have
higher 206Pb/204Pb than the Kita-Matsuura basalts, but
similar 208Pb/204Pb to them.
The isotopic compositions of the Kita-Matsuura basalts
show systematic spatial and temporal variations (Fig. 7b,
d, and f). Samples from the easternmost Kunimi section
(Figs 1 and 7g), have relatively depleted Sr^Nd isotope
compositions, and higher 208Pb/204Pb for a given
206
Pb/204Pb. The Kunimi section lavas become more
depleted MORB mantle (DMM)-like with time (Zindler
& Hart, 1986). Samples from the Ishimori section (Figs 1
and 7g) plot in the middle of the range of the
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MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Fig. 4. Relationships between Nb concentration and (a) Nb/Ce and (b) Nb/Th. Variations of Ce/Yb (c) and Sm/Yb (d) are also shown for the
three basalt groups as a function of eruption age. In (a) and (b) the symbols define the different basalt groups. In (c) and (d) data from the
same section are connected by lines.
Kita-Matsuura basalts in Sr^Nd^Pb isotope space. The
Ishimori section temporally changes to higher 206Pb/204Pb
(Fig. 7d) and 207Pb/204Pb. The three early samples from
the Senryu section (Figs 1 and 7g) temporally approach a
depleted component in terms of their Sr^Nd^Pb composition, but the three later samples temporally change towards an enriched component in Sr^Nd space and a
component with higher 206Pb/204Pb, 207Pb/204Pb, and
208
Pb/204Pb values. Samples from the Hirado section
(Figs 1 and 7g), the westernmost section, plot in the most
enriched region for the Kita-Matsuura basalts in Sr^Nd
space. They show a temporal change to higher
206
Pb/204Pb (Fig. 7d), 207Pb/204Pb, and 208Pb/204Pb
(Fig. 7d).
There is also a systematic correlation between Pb isotope composition and the SiO2 groups: 206Pb/204Pb for
given values of 143Nd/144Nd and 208Pb/204Pb increases in
the order of the low-, medium-, and high-SiO2 groups
(Fig. 7d). The range of the low-SiO2 group does not overlap
that of the high-SiO2 group in the Pb isotope system,
whereas the medium-SiO2 group lies between the lowand high-SiO2 groups (Fig. 7d). The 87Sr/86Sr,
143
Nd/144Nd, and 206Pb/204Pb of the low-SiO2 samples in
the earliest stages of the Senryu, Ishimori, and Kunimi
sections are intermediate within the range of variation of
the Kita-Matsuura samples (Fig. 7c, d, and f). Samples
from the high-SiO2 group, especially in the latest stages of
the Hirado, Senryu, and Ishimori sections, have the highest 206Pb/204Pb among the Kita-Matsuura basalts, except
for sample 02110934, which has extremely high
208
Pb/204Pb (Fig. 7c, d, and f).
E S T I M AT I O N O F M E LT I N G
CON DITIONS
In the following section, the three least-fractionated basalts
are used to represent the continuous spectrum of parental
magmas from which the entire chemical diversity of
the Kita-Matsuura basalts is derived (Sakuyama et al.,
2009). These are referred to as parental magma 1
[PM1 (¼ 031027e4) hereafter] for the low-SiO2 group,
PM2 (¼ 02111173b) for the medium-SiO2 group, and
PM3 (¼ 02110934) for the high-SiO2 group.
The estimation of melting conditions was performed in
four steps. First, the compositions of the primary melts in
equilibrium with mantle peridotite were estimated from
PM1, PM2, and PM3 by adding olivine (Sakuyama et al.,
2009), following Tamura et al. (2000). Addition of clinopyroxene with olivine was investigated to evaluate how much
this can affect the results; this did not have any significant
effect on the results in the following discussion. The detailed procedure has been described by Sakuyama et al.
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Table 2: Representative chemical compositions of spinel inclusions
Host rock:
031027E4
031027E4
031027E4
031027E4
031027E4
031027E4
031027E4
031027E4
Basalt group:
low
low
low
low
low
low
low
low
Section:
Kunimi
Kunimi
Kunimi
Kunimi
Kunimi
Kunimi
Kunimi
Kunimi
Sample:
sp1–ol6
sp1–ol22
sp1–ol55
sp1–ol30
sp2–ol2
sp1–ol67
sp2–ol81
sp1–ol86
Major elements (wt %)
SiO2
0·05
0·12
Al2O3
14·13
17·96
0·07
14·0
0·08
0·07
0·06
0·02
0·07
19·42
20·73
15·32
16·84
18·04
TiO2
0·75
0·86
0·62
1·05
0·86
0·76
0·95
0·63
FeO*
37·20
25·51
36·11
25·82
23·59
33·04
30·85
27·05
MnO
0·42
0·26
0·39
0·22
0·25
0·32
0·26
0·29
MgO
6·04
10·45
6·59
10·86
12·07
7·14
8·93
9·61
CaO
0·03
0·05
0·04
0·01
0·02
0·00
0·00
0·01
Na2O
0·00
0·01
0·00
0·03
0·00
0·02
0·01
0·01
K2O
0·00
0·01
0·01
0·02
0·01
0·01
0·00
0·00
Cr2O3
38·72
44·03
38·49
41·68
42·37
41·27
40·34
42·48
0·18
V2O3
0·19
0·15
0·22
0·12
0·22
0·21
0·11
NiO
0·14
0·12
0·09
0·13
0·14
0·05
0·09
0·09
Total
97·67
99·54
97·42
99·42
100·31
98·19
98·39
98·44
Cations
Si
0·010
0·023
0·016
0·015
0·013
0·012
0·005
0·013
Al
3·538
4·132
3·684
4·440
4·629
3·732
4·015
4·223
Ti
0·119
0·127
0·098
0·153
0·122
0·118
0·144
0·094
Fe2þ
6·611
4·163
6·380
4·188
3·738
5·710
5·220
4·493
Mn
0·075
0·044
0·069
0·036
0·041
0·055
0·044
0·049
Mg
1·914
3·040
2·075
3·139
3·410
2·200
2·693
2·844
Ca
0·007
0·011
0·009
0·003
0·004
0·000
0·000
0·001
Na
0·000
0·005
0·000
0·009
0·000
0·006
0·005
0·002
K
0·000
0·002
0·004
0·004
0·001
0·001
0·000
0·001
Cr
6·503
6·792
6·428
6·391
6·348
6·742
6·453
6·671
V
0·033
0·024
0·037
0·019
0·034
0·035
0·017
0·028
Ni
0·024
0·019
0·015
0·020
0·021
0·008
0·015
0·014
Cation total
1·834
18·380
18·814
18·415
18·360
18·619
18·611
18·434
100Cr3þ/(Cr3þ þ Al3þ)
64·77
62·18
63·57
59·01
57·83
64·37
61·64
61·24
100Cr3þ/(Cr3þ þ Al3þ þ Fe3þ)
48·88
54·72
48·27
51·38
51·22
52·22
50·19
52·96
Host olivine Fo#
77·37
87·23
80·17
86·34
87·88
78·06
82·53
84·74
Host rock:
031027E4
02111173
02111173
02111173
02111173
02111173
02110934
02110934
Basalt group:
low
medium
medium
medium
medium
medium
high
high
Section:
Kunimi
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Sample:
sp1–ol89
sp1–ol4
sp2–ol4
sp1–ol10
sp1–ol15
sp1–ol16
sp4–ol13
sp2–0l10
Major elements (wt %)
SiO2
0·03
0·06
0·06
0·03
0·04
0·03
0·10
0·10
Al2O3
15·10
14·62
15·86
18·17
9·48
14·39
22·19
25·64
TiO2
0·91
1·34
1·46
1·24
1·31
1·14
1·19
1·16
FeO*
29·45
39·04
39·29
33·03
47·31
42·21
26·57
26·69
MnO
0·33
0·42
0·39
0·38
0·45
0·38
0·21
0·23
MgO
8·43
6·10
6·51
8·24
4·70
5·91
10·66
11·71
(continued)
1092
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Table 2: Continued
Host rock:
031027E4
02111173
02111173
02111173
02111173
02111173
02110934
02110934
Basalt group:
low
medium
medium
medium
medium
medium
high
high
Section:
Kunimi
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Sample:
sp1–ol89
sp1–ol4
sp2–ol4
sp1–ol10
sp1–ol15
sp1–ol16
sp4–ol13
sp2–0l10
CaO
0·01
0·06
0·02
0·02
0·02
0·07
0·01
0·03
Na2O
0·00
0·02
0·03
0·01
0·00
0·00
0·01
0·00
0·00
0·01
0·00
0·01
0·01
0·01
0·01
0·00
Cr2O3
K2O
43·86
33·27
31·90
36·18
32·42
30·78
37·46
31·68
V2O3
0·19
0·18
0·17
0·12
0·20
0·18
0·21
0·16
NiO
0·07
0·12
0·12
0·11
0·14
0·11
0·15
0·16
Total
98·36
95·22
95·81
97·52
96·08
95·20
98·76
97·57
Cations
Si
0·006
0·013
0·012
0·005
0·009
0·006
0·020
0·020
Al
3·622
3·762
4·027
4·377
2·565
3·751
5·048
5·794
Ti
0·139
0·220
0·237
0·190
0·226
0·189
0·173
0·167
Fe2þ
5·013
7·128
7·077
5·647
9·080
7·806
4·289
4·279
Mn
0·056
0·078
0·072
0·065
0·087
0·071
0·034
0·037
Mg
2·558
1·984
2·090
2·510
1·606
1·948
3·067
3·347
Ca
0·003
0·013
0·005
0·004
0·006
0·017
0·001
0·006
Na
0·000
0·007
0·013
0·003
0·000
0·000
0·004
0·000
K
0·000
0·002
0·001
0·003
0·004
0·002
0·001
0·001
Cr
7·059
5·743
5·432
5·846
5·881
5·382
5·718
4·802
V
0·030
0·031
0·028
0·19
0·037
0·032
0·033
0·025
Ni
0·011
0·021
0·021
0·018
0·025
0·019
0·023
0·024
Cation total
18·449
19·003
19·015
18·687
19·525
19·224
18·410
18·503
100Cr3þ/(Cr3þ þ Al3þ)
66·06
60·42
57·43
57·19
69·63
58·93
53·11
45·32
100Cr3þ/(Cr3þ þ Al3þ þ Fe3þ)
55·8
42·78
40·47
45·36
40·68
38·60
46·14
38·14
Host olivine Fo#
82·01
78·99
80·32
82·91
75·16
78·27
85·97
85·56
Host rock:
02110934
02110934
02110934
02110934
02110934
02110934
Basalt group:
high
high
high
high
high
high
02110934
high
Section:
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Sample:
sp1–ol18
sp3–ol8
sp1–ol16
sp1–ol30
sp1–ol21
sp1–ol24
sp1–ol20
Major elements (wt %)
SiO2
0·05
0·08
0·09
0·08
0·06
0·07
0·08
Al2O3
20·83
24·55
18·25
19·89
21·76
23·52
19·33
TiO2
1·36
0·71
0·67
0·91
0·79
0·90
1·02
FeO*
29·22
20·55
24·94
26·42
24·51
25·84
26·51
MnO
0·33
0·21
0·29
0·26
0·26
0·31
0·28
MgO
9·70
11·82
10·31
10·51
12·06
11·61
10·57
CaO
0·02
0·00
0·03
0·00
0·00
0·00
0·09
Na2O
0·00
0·02
0·00
0·00
0·00
0·00
0·00
K2O
0·00
0·01
0·01
0·00
0·01
0·00
0·00
Cr2O3
36·03
3·98
42·21
37·89
38·55
33·58
39·36
0·17
V2O3
0·24
0·12
0·09
0·13
0·16
0·14
NiO
0·10
0·19
0·16
0·14
0·18
0·18
0·11
Total
97·87
97·23
97·05
96·23
98·34
96·15
97·52
(continued)
1093
JOURNAL OF PETROLOGY
VOLUME 55
NUMBER 6
JUNE 2014
Table 2: Continued
Host rock:
02110934
02110934
02110934
02110934
02110934
02110934
02110934
Basalt group:
high
high
high
high
high
high
high
Section:
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Sample:
sp1–ol18
sp3–ol8
sp1–ol16
sp1–ol30
sp1–ol21
sp1–ol24
sp1–ol20
Si
0·010
0·015
0·019
0·015
0·011
0·014
0·016
Al
4·859
5·505
4·292
4·695
4·943
5·436
4·515
Ti
0·203
0·101
0·101
0·136
0·114
0·133
0·153
Fe2þ
4·837
3·270
4·162
4·424
3·951
4·237
4·393
Mn
0·055
0·034
0·049
0·044
0·043
0·051
0·047
Mg
2·860
3·352
3·068
3·137
3·464
3·394
3·123
Ca
0·003
0·000
0·006
0·000
0·001
0·001
0·019
Na
0·000
0·008
0·000
0·000
0·000
0·000
0·000
K
0·001
0·001
0·001
0·000
0·003
0·000
0·000
Cr
5·638
5·864
6·659
5·997
5·874
5·206
6·167
V
0·037
0·018
0·015
0·021
0·025
0·022
0·026
Ni
0·016
0·029
0·026
0·023
0·028
0·028
0·018
Cation total
18·520
18·196
18·398
18·492
18·456
18·521
18·477
100Cr3þ/(Cr3þ þ Al3þ)
53·71
51·58
60·81
56·09
54·30
48·92
57·73
100Cr3þ/(Cr3þ þ Al3þ þ Fe3þ)
44·96
48·44
53·26
47·47
46·64
41·04
49·13
Host olivine Fo#
84·54
87·77
86·95
86·02
87·18
86·30
86·84
Cations
*Total iron given as FeO.
Cation numbers are calculated on the basis of 24 O. Ferric content for spinel was estimated by assuming spinel stoichiometry All Ti was combined with Fe as the ulvöspinel component (Fe2TiO4) in the calculation. All analysed spinel grains
are inclusions in olivine phenocrysts.
(2009). Second, the fertility of the residual source peridotite
was estimated based on the compositional relationship between the spinel inclusions and the host olivine phenocrysts. Third, water contents in the primary melts were
estimated using plagioclase^melt and olivine^melt hygrometers and the observed crystallization sequence of the
basalts. Fourth, melting pressures and temperatures were
estimated under hydrous conditions by using the estimated
primary magma compositions and published experimental
data. Details of the procedures employed are presented in
the following sections.
Estimation of primary melt compositions
Arai (1987) proposed the use of the olivine^spinel mantle
array (OSMA) as an indicator of the degree of melt extraction from a peridotite source on the basis of a positive
correlation between the Fo# of olivine and Cr# of chromian spinel in peridotite. Once a primary melt segregates
from the source peridotite and starts fractionating olivine
phenocrysts, the Fo# of olivine in equilibrium with the residual melt quickly decreases as olivine fractionation progresses; this creates a fractionation trend at a high angle
to the OSMA. Because Cr# of both the low- and
high-SiO2 groups is 50 at Fo# of 88, olivine in equilibrium with the primary melts of the Kita-Matsuura basalt
should have a relatively high Fo#. If the peridotite source
mantle beneath northwestern Kyushu follows the OSMA,
the Fo# of residual olivine in equilibrium with the primary melt of the Kita-Matsuura basalt is estimated to be
90 or more by linear extrapolation of the observed trend
(Fig. 5a). Here, we assume Fo# 90 for olivine in equilibrium with the primary melt of the low-SiO2 group.
The melting degree of the source mantle for the lowSiO2 group is estimated to be the lowest of the two groups
of the Kita-Matsuura basalt (Sakuyama et al., 2009). The
corresponding Fo# for the medium- and high-SiO2
groups are thus inferred to be greater than 90, if they
were all derived from the same original source mantle.
The rate of increase of Cr# in spinel relative to the increase of Fo# in the host olivine increases with the progress of melting, as shown by the OSMA. The estimated
Fo# of olivine in the source peridotite for the high-SiO2
group also has a similar composition to the low-SiO2
group, judging from the intercept at Cr# ¼ 55^60 on the
OSMA (Fig. 5a). This suggests a similar degree of depletion of the sources of the Kita-Matsuura primary basalts
1094
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
the partial melt (e.g. Hirose & Kawamoto, 1995). It is thus
very important to evaluate the effect of water when
estimating melting conditions. Here, we examine water
content in the primary melt for each SiO2 group by combining two approaches. First, we estimated water contents
in the melt using a combination of the plagioclase^liquid
hygrometer of Lange et al. (2009) and a geothermometer
for olivine-saturated melts (Sugawara, 2000; Medard &
Grove, 2008). This hybrid method utilizes the difference of
liquidus temperature decrease between plagioclase and
olivine when water is added to the melt. Second, we modelled crystallization trends using MELTS (Ghiorso &
Sack, 1995) to test whether the water contents estimated
by the geohygrometer can reasonably reproduce the geochemical trend of each SiO2 group.
Crystal^liquid hygrometer^thermometer
Lange et al. (2009) developed a new thermodynamic model
for the plagioclase^liquid exchange reaction between
albite and anorthite components, which can be used as a
plagioclase^liquid hygrometer. The hygrometer is formulated as
wt % H2 O ¼ m0 f ðP,TÞ þ a00 þ
Fig. 5. (a) Relationships between the Cr/(Cr þAl) atomic ratio
(Cr#) of spinel inclusions and the Fo content (Fo#) of the host olivine in the least fractionated olivine basalt (02110934 ¼ PM3, highSiO2, Senryu) and olivine^clinopyroxene basalt (031027e4 ¼ PM1,
low-SiO2, Kunimi) from the Kita-Matsuura basalts. Continuous-line
and dashed-line arrows represent fractionation trends defined by clusters of data. (b) Cr^Al^Fe3þ ternary diagram showing the composition of chromian spinel included in olivine phenocrysts within the
two least fractionated basalts.
and contradicts the previous assumptions (Fo# ¼ 89^91)
of Sakuyama et al. (2009). Therefore, we first assumed that
the Fo# of olivine in equilibrium with the primary melt
of the medium- and high-SiO2 groups was also 90 (the
effect of a residue with Fo# higher than 90 will be discussed below). The amount of olivine or olivine þ clinopyroxene added to PM1, PM2, and PM3 was 13, 18, and 16 wt
% for olivine addition and 18, 24, and 21wt % for olivine þ clinopyroxene addition, respectively. These values
were taken from Sakuyama et al. (2009). The estimated
MgO contents in the primary melts are 16·6, 16·2, and
15·3 wt % and FeO* contents are 10·5, 10·4, and 9·9 wt %
when olivine was added for the low-, medium-, and highSiO2 groups, respectively (Table 1).
Estimation of water content of the magma
b00 X 00
þ
di X i
T
ð1Þ
where m0, a00, b00, and d00 i are constant parameters that were
estimated by fitting during calibration of experimental
data, P is the pressure, T the temperature, and Xi is the
fraction of the oxide component i in the melt. The water
content in a melt is a function of the chemical compositions
of the plagioclase and melt, temperature, and pressure.
Activity^composition relationships for the plagioclase
solid solution were taken from the Holland & Powell
(1992) model 4. When we applied this model to natural
samples it was necessary to constrain the crystallization
pressure and temperature. When the pressure was fixed,
the water content could be expressed as a function of
temperature.
If the plagioclase-saturated melt is also saturated with
olivine, the olivine liquidus temperature can be estimated
by applying the melt geothermometer equation for an olivine-saturated liquid proposed by Sugawara (2000). The
thermometer is formulated as
Liq
Liq
TðKÞ ¼ 1446 1 440XSiO2 0 5XFeO
Liq
Liq
ð2Þ
þ 12 32XMgO 3 899XCaO þ 0 0043P
where X is mol % and P is the pressure (0·1MPa). The
effect of water on the olivine liquidus temperature was
evaluated by applying the thermodynamic model proposed
by Medard & Grove (2008). This model equates to
Water in the source peridotite, even in small amounts, can
affect the degree of melting at a given temperature (e.g.
Green, 1973). Accordingly, this affects the composition of
1095
melt
melt 2
melt 3
2 97ðCH
Þ þ 0 0761ðCH
Þ
TðKÞ ¼ 40 4CH
2O
2O
2O
ð3Þ
JOURNAL OF PETROLOGY
VOLUME 55
NUMBER 6
JUNE 2014
Table 3: Representative chemical compositions of plagioclase phenocrysts or microphenocrysts
Host rock:
031026A5
031026A5
031026A5
03102045
03102045
03102045
03102147
03102147
03102147
03102151
Basalt group:
low
low
low
low
low
low
medium
medium
medium
medium
Section:
Kunimi
Kunimi
Kunimi
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Senryu
Sample:
pl-1
pl-2
pl-3
pl-1
pl-2
pl-3
pl-1
pl-2
pl-3
pl-1
SiO2
51·19
52·49
55·10
47·37
49·65
52·13
50·07
51·90
54·51
51·89
Al2O3
30·16
29·04
27·70
32·12
30·08
28·85
31·11
29·63
27·56
29·38
TiO2
0·07
0·09
0·04
0·06
0·06
0·13
0·06
0·0
0·14
0·08
FeO*
0·35
0·45
0·35
0·41
0·46
0·72
0·42
0·35
1·06
0·51
MnO
0·00
0·02
0·00
0·04
0·04
0·00
0·00
0·02
0·04
0·01
MgO
0·08
0·08
0·08
0·12
0·08
0·13
0·09
0·13
0·16
0·09
CaO
13·52
12·37
11·04
15·87
14·08
12·33
14·52
12·90
11·05
13·39
Na2O
3·66
4·27
5·07
2·28
3·23
4·23
3·18
3·83
4·96
3·90
K2O
0·26
0·32
0·40
0·15
0·27
0·42
0·15
0·18
0·38
0·30
Cr2O3
0·00
0·00
0·02
0·02
0·00
0·01
0·03
0·00
0·00
0·00
V2O3
0·02
0·00
0·01
0·00
0·04
0·01
0·00
0·00
0·00
0·00
NiO
0·00
0·00
0·00
0·01
0·01
0·00
0·00
0·01
0·05
0·00
Total
99·30
99·13
99·82
98·44
98·00
98·96
99·62
99·00
99·90
99·55
Si
7·042
7·216
7·484
6·629
6·946
7·196
6·887
7·141
7·432
7·127
Al
4·892
4·706
4·436
5·298
4·960
4·695
5·044
4·806
4·429
4·757
Ti
0·007
0·009
0·004
0·006
0·006
0·014
0·007
0·005
0·014
0·008
Fe2þ
0·040
0·052
0·039
0·049
0·054
0·083
0·048
0·040
0·121
0·058
Mn
0·000
0·002
0·000
0·004
0·005
0·000
0·000
0·003
0·004
0·001
Mg
0·017
0·016
0·017
0·026
0·017
0·027
0·018
0·026
0·032
0·019
Ca
1·992
1·823
1·607
2·379
2·111
1·824
2·140
1·902
1·615
1·971
Na
0·977
1·137
1·336
0·620
0·876
1·132
0·848
1·022
1·312
1·037
K
0·045
0·056
0·070
0·026
0·048
0·074
0·027
0·031
0·066
0·053
Cr
0·000
0·000
0·002
0·002
0·000
0·001
0·003
0·000
0·000
0·000
V
0·002
0·000
0·001
0·000
0·004
0·001
0·000
0·000
0·000
0·000
Ni
0·000
0·000
0·000
0·001
0·001
0·000
0·000
0·002
0·005
0·000
Cation total
15·015
15·019
14·995
15·038
15·028
15·045
15·021
14·978
15·029
15·031
An#
66·09
60·43
53·34
79·34
70·68
61·71
71·63
65·06
55·17
65·52
KDðCa=NaÞ
1·18
0·93
0·70
1·49
0·94
0·63
1·38
1·02
0·67
1·29
Host rock:
03102151
03102151
02110607
02110607
02110607
03101923
03101923
Basalt group:
medium
medium
high
high
high
high
high
high
Section:
Senryu
Senryu
Senryu
Senryu
Senryu
Hirado
Hirado
Hirado
Sample:
pl-2
pl-3
pl-1
pl-2
pl-3
pl-1
pl-2
pl-3
03101923
SiO2
53·17
56·35
50·88
53·19
56·40
49·86
51·04
51·07
Al2O3
28·84
26·91
30·17
28·83
2·71
31·09
30·19
30·41
TiO2
0·06
0·00
0·01
0·10
0·11
0·05
0·06
0·06
FeO*
0·27
0·25
0·75
0·83
0·68
0·64
0·69
0·62
MnO
0·03
0·00
0·00
0·01
0·01
0·00
0·01
0·00
MgO
0·06
0·03
0·05
0·06
0·04
0·06
0·04
0·06
CaO
12·41
10·07
13·47
11·92
9·66
14·69
13·88
13·39
Na2O
4·56
5·72
3·54
4·40
5·72
3·06
3·38
3·61
K2O
0·31
0·45
0·21
0·27
0·45
0·20
0·22
0·23
Cr2O3
0·00
0·00
0·01
0·00
0·02
0·00
0·00
0·00
(continued)
1096
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Table 3: Continued
Host rock:
03102151
03102151
02110607
02110607
02110607
03101923
03101923
Basalt group:
medium
medium
high
high
high
high
high
03101923
high
Section:
Senryu
Senryu
Senryu
Senryu
Senryu
Hirado
Hirado
Hirado
Sample:
pl-2
pl-3
pl-1
pl-2
pl-3
pl-1
pl-2
pl-3
V2O3
0·00
0·00
0·05
0·02
0·05
0·00
0·01
0·03
NiO
0·00
0·00
0·01
0·03
0·00
0·00
0·00
0·00
Total
99·69
99·79
99·15
99·64
99·84
99·65
99·50
99·46
Cations
Si
7·265
7·634
7·022
7·272
7·644
6·869
7·020
7·019
Al
4·644
4·298
4·907
4·647
4·267
5·048
4·895
4·927
Ti
0·006
0·000
0·001
0·010
0·011
0·005
0·006
0·006
Fe2þ
0·030
0·029
0·086
0·095
0·077
0·073
0·080
0·071
Mn
0·003
0·000
0·000
0·001
0·001
0·000
0·001
0·000
Mg
0·012
0·006
0·010
0·013
0·009
0·012
0·008
0·012
Ca
1·817
1·461
1·992
1·746
1·404
2·169
2·046
1·972
Na
1·209
1·503
0·947
1·165
1·503
0·818
0·901
0·961
K
0·053
0·078
0·03
0·046
0·078
0·036
0·039
0·040
Cr
0·000
0·000
0·001
0·000
0·002
0·000
0·000
0·000
V
0·000
0·000
0·005
0·002
0·005
0·000
0·000
0·000
0·000
0·000
0·001
0·003
0·000
0·000
0·000
0·000
Cation total
Ni
15·038
15·008
15·011
15·000
14·999
15·029
14·996
15·010
An#
60·05
49·29
66·94
59·03
47·03
71·77
68·53
66·34
1·02
0·66
1·27
0·90
0·56
1·57
1·34
1·21
KDðCa=NaÞ
*Total iron given as FeO.
Cation numbers were calculated on the basis of 24 oxygen atoms. KDðCa=NaÞ was calculated by assuming a whole-rock
composition as the melt composition in equilibrium with the plagioclase.
melt
where CH
is the wt % of H2O in the melt. Medard &
2O
Grove (2008) experimentally quantified the effects of
water on the liquidus of olivine-saturated primitive basaltic
and andesitic melts. They estimated Tat a given pressure
with a known amount of water added, and successfully
separated the effect of water from other potential influences (e.g. melt composition and pressure) over a wide
range of olivine-saturated basaltic compositions. Although
we do not have a precise constraint on the crystallization
pressure, clinopyroxene phenocryst compositions suggest
that the crystallization pressure was less than 0·5 GPa
(Sakuyama et al., 2009). Therefore, we varied pressure conditions from 0·1 to 0·5 GPa in this study.
In contrast to plagioclase, the liquidus temperature drop
of olivine is less sensitive to the water content in the melt
than that of plagioclase. The temperature^water content
in melt relationship for plagioclase and olivine should
cross over at a specific temperature and water content,
where olivine and plagioclase are both saturated under
the same conditions (Fig. 8a).
We estimated the water content in the melt from the
intersection of the Lange et al. (2009) plagioclase^liquid
hygrometer and the olivine-saturated melt geothermometer^hygrometer of Sugawara (2000) and Medard &
Grove (2008) (Fig. 8a). For samples in which both olivine
and plagioclase are saturated, and equilibrium among olivine, plagioclase, and melt can be assumed, we can estimate
the temperature and water content of the melt at a given
pressure. To evaluate the accuracy of this approach, we
applied this method to the results of experimental studies
from the literature under hydrous conditions. Standard deviations (1s) of the difference of the temperature and
water content between the estimated melt and experiments
were 238C and 0·6 wt %, respectively (see Appendix for
details).
To ensure equilibrium, we selected aphyric or nearly
aphyric samples that contain euhedral olivine and plagioclase phenocrysts, which can be assumed to be in equilibrium with a melt close to the whole-rock composition. We
assumed that the phenocryst cores with the highest Fo#
and An# in the samples were in equilibrium with the
melt.
Figure 8b and Table 5 shows the estimated water contents in the melt assuming 0·3 GPa. The error bars
1097
JOURNAL OF PETROLOGY
VOLUME 55
NUMBER 6
JUNE 2014
respectively. In contrast, the minimum values of the estimated water contents in the samples of the low-,
medium-, and high-SiO2 groups were 0·15, 0·26, and
0·35 wt %, respectively. These low values of H2O may reflect crystallization of the plagioclase phenocrysts after
degassing during magma ascent to the surface. Because
degassing of H2O from the magma before crystallization
of plagioclase may have occurred in all the samples, the
maximum water contents estimated above are minimum
estimates of the water content of the magma.
Our estimate of the water content of the primary melts is
slightly lower than that of Iwamori (1991); his estimated
water contents in alkaline and tholeiitic basalts of the
Chugoku district, southwestern Japan, were 0·5^1·5 wt %
and 1wt %, respectively. These values were obtained
from a melting experiment. Iwamori (1991) also considered
the effect of CO2. However, the contribution of CO2 to
the generation of the Kita-Matsuura basalt may be
limited because the Kita-Matsuura basalt is not as silicaundersaturated as the alkaline basalt in Chugoku.
Estimation by MELTS calculations
Fig. 6. Histograms of anorthite (An) content of the cores of plagioclase phenocrysts or microphenocrysts for (a) low-SiO2 (b) mediumSiO2 and (c) high-SiO2 basalt groups. Microphenocrysts were analyzed in 03101923 because of the absence of phenocrysts. In (a) and
(b) arrows indicate the compositional range of groundmass plagioclase, whereas in (c) they indicate the compositions of the outermost
rim of plagioclase microphenocrysts.
represent the range of water contents obtained by varying
pressure from 0·1 to 0·5 GPa. Because the samples show
variable extents of differentiation (Fig. 8b), the water content in the primary melt was then calculated from the fractionation-corrected least differentiated samples (PM1,
PM2, and PM3) of each SiO2 group. The maximum estimated water contents in the low-, medium-, and highSiO2 groups are 0·7, 1·3, and 2·1wt %, respectively.
Combined with the amount of olivine fractionated from
the PM1^PM3, the maximum water contents in the least
differentiated melts of low-, medium-, and high-SiO2
groups were estimated to be 0·6, 1·1, and 1·8 wt %,
To cross-check the water contents estimated above, we performed MELTS calculations to investigate whether the
major element trends for the low-, medium-, and highSiO2 groups can be reproduced by isobaric fractional and
equilibrium crystallization starting from parental
magmas PM1, PM2, and PM3 (Fig. 9a^c). Water contents
assumed in the initial melts were 0·25, 0·5, 1 and 2 wt %,
and crystallization pressures were set to 0·1, 0·2 and
0·3 GPa. Isobaric crystallization paths were calculated
with increments of 2 K decrease in temperature under
oxygen fugacity conditions 1 log unit less than the Ni^
NiO oxygen fugacity buffer (NNO 1) (Nilsson & Peach,
1993; Farley, 1994); the results were then compared with
the observed fractionation trends.
Although both the fractional and equilibrium crystallization paths are similar, fractional crystallization tends
to generate higher TiO2, FeO*, Na2O, K2O, and P2O5,
and lower CaO, MgO, and Al2O3. The crystallization
paths calculated by MELTS for TiO2, Al2O3, and FeO*
during both fractional and equilibrium crystallization are
more sensitive to water content than to crystallization pressure (Fig. 9d^f, g^i, m^o). The TiO2 content increases
with decreasing water content of the initial melt, particularly during the early crystallization stage (Fig. 9j and m).
This behaviour of TiO2 is mainly controlled by the liquidus temperature of plagioclase and Fe^Ti oxides, which is
strongly dependent on the water content (Housh & Luhr,
1991; Takagi et al., 2005). As water content in the melt increases, the difference of liquidus temperature between
plagioclase and Fe^Ti oxides decreases, whereas the liquidus temperatures of both plagioclase and Fe^Ti oxides
decrease. If the melt is enriched in water, the amount
of fractionation of plagioclase before the onset of
1098
03101930
Sample:
1099
0·006
38·660
0·002
15·584
0·003
18·378
0·000012
0·512698
0·000013
0·003
38·565
0·001
15·550
0·001
18·296
0·000011
0·512688
0·000015
0·704638
03101924
medium
8·1
Hirado
0·008
38·483
0·003
15·562
0·004
18·286
0·000012
0·512574
0·000013
0·705479
03101817
medium
8·84
Hirado
0·012
38·749
0·005
15·603
0·006
18·353
0·000013
0·512724
0·000016
0·70441
02110723
high
6·37
Senryu
0·018
39·125
0·007
15·715
0·008
18·451
0·000013
0·51277
0·000017
0·704298
02110934
high
6·48
Senryu
0·007
38·506
0·003
15·528
0·003
18·272
0·000010
0·512765
0·000012
0704319
02111276
medium
6·54
Senryu
0·014
38·461
0·006
15·514
0·006
18·209
0·000009
0·512796
0·000013
0·703647
02111173B
medium
6·56
Senryu
0·012
38·637
0·005
15·541
0·006
18·237
0·000008
0·512801
0·000015
0·704067
03102155
medium
6·59
Senryu
0·015
38·784
0·006
15·537
0·007
18·248
0·000010
0·512751
0·000013
0·704888
03102145
low
7·68
Senryu
0·003
38·691
0·001
15·576
0·001
18·374
0·000007
0·512744
0·000015
0·704247
031029if7
high
5·42
Ishimori
0·003
38·616
0·001
15·545
0·002
18·289
0·000010
0·51271
0·000014
0·704384
03102481
high
7·33
Ishimori
0·004
38·881
0·002
15·533
0·002
18·216
0·000012
0·512715
0·000013
0·704117
03102475
low
7·74
Ishimori
0·013
38·326
0·005
15·493
0·006
18·036
0·000010
0·512866
0·000012
0·70362
031027e4
low
7·11
Kunimi
0·018
38·413
0·007
15·552
0·009
18·253
0·000010
0·512757
0·000012
0·703845
031027c9
low
7·26
Kunimi
0·003
38·917
0·001
15·565
0·002
18·292
0·000010
0·512697
0·000014
0·704552
031026a8
low
7·76
Kunimi
K–Ar ages were calculated by regressing separately for each section to estimate ages for lava flows without age determination [see details given by Sakuyama et al.
(2009)].
1
2s
Pb/204Pb
208
2s
Pb/204Pb
207
2s
Pb/204Pb
206
2s
Nd/144Nd
143
2s
0·704973
high
SiO2 group:
Sr/86Sr
7·29
K–Ar age (Ma)1:
87
Hirado
Section:
Table 4: Sr^Nd^Pb isotopic compositions of the Kita-Matsuura basalt
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
JOURNAL OF PETROLOGY
VOLUME 55
NUMBER 6
JUNE 2014
Fig. 7. Sr^Nd^Pb isotopic compositions of the Kita-Matsuura basalts. (a, b) 143Nd/144Nd vs 87Sr/86Sr; (c, d) 208Pb/204Pb vs 206Pb/204Pb; (e, f)
143
Nd/144Nd vs 206Pb/204Pb; (b), (d), and (f) are close-ups of the rectangles in (a), (c), and (e), respectively. Numbers in symbols in (b), (d),
and (f) indicate the eruption sequence as shown in (g). AOC, altered oceanic crust. The northern hemisphere reference line (NHRL; Hart,
1984) and magnitude of deviation from the NHRL (8/4 and 7/4) are shown as dashed^dotted lines in (c). Filled inverted triangle represents
a 15 Ma dacite sample reported by Uto et al. (2004). Isotopic end-member components for DMM, EM1, EM2, and HIMU are from Zindler &
Hart (1986) and Hofmann (1997). PAC, Pacific Plate; PHS, Philippine Sea Plate.
1100
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
1^2 wt %, respectively. Within these ranges of water contents, a crystallization pressure of 0·2 GPa was predicted
for every group. Differentiation paths calculated for pressures greater than 0·2 GPa, assuming both fractional and
equilibrium crystallization, predict a decrease in SiO2
during the early stages of crystallization owing to excessive
fractionation of clinopyroxene, with the extent of the SiO2
decrease increasing with pressure. Crystal fractionation at
pressures higher than 0·2 GPa is, therefore, unlikely as suggested by Sakuyama et al. (2009).
The estimates of melt H2O contents based on MELTS
calculations are consistent with those estimated using a
combination of the plagioclase^liquid hygrometer and olivine-saturated liquid hygrometer. We adopted the maximum
values of the H2O contents estimated by the hygrometers
as the H2O contents in the least-differentiated samples of
each SiO2 group: the values are 0·6 wt %, 1·1wt %, and
1·8 wt % for the low-, medium-, and high-SiO2 groups, respectively. Taking the amount of crystal fractionation from
the primary melts into account, the H2O contents in the primary melt of each SiO2 group were estimated to be 0·5 wt
%, 0·9 wt %, and 1·5 wt %, respectively.
Estimation of melting conditions
Melting pressure and the effect of H2O
Fig. 8. (a) Relationships between water content estimated using the
plagioclase hygrometer of Lange et al. (2009) (dashed and dotted
curves) and the olivine-saturated melt thermometer (Sugawara,
2000; Medard & Grove, 2008) (continuous curves) for the mediumSiO2 sample 03102147 from the Senryu section. The intersection of
the dashed and continuous curves gives the water content in the melt
and crystallization temperature. (b) Relationship between FeO*/
MgO and water content in the melt estimated assuming a pressure of
0·3 GPa. Error bars indicate the range of water contents for pressures
from 0·1 to 0·5 GPa. (c) Comparison of estimated water content in
the melt for the three basalt groups. Range of water content in the
melt estimated by MELTS is indicated by a double-headed arrow.
Fe^Ti oxide fractionation decreases. In this case, TiO2 content of the melt does not increase with decreasing MgO.
Tatsumi & Suzuki (2009) showed that increased TiO2 content at a given SiO2 content for less hydrous melts was
caused by greater fractionation of plagioclase (and lesser
fractionation of Fe^Ti oxides). Their conclusion is also consistent with our observations.
Plausible ranges of water content in the PM1^PM3
magmas are considered to be 0·25^0·5, 0·25^1, and
Although the Mg/Si ratio of a partial melt of peridotite decreases with increasing water content in the system (e.g.
Hirose, 1997), the decrease in Mg/Si ratio is suppressed
with increasing pressure up to 5 GPa (Kushiro et al., 1968;
Inoue & Sawamoto, 1992; Inoue, 1994). In melting experiments in the water-bearing peridotite system at 1GPa,
Hirose & Kawamoto (1995) showed that partial melts
with less than 2·5 wt % H2O, which formed at temperatures above 12008C, are all within the range of major element compositions of anhydrous melts formed by the same
degree of partial melting under pressures from 0·5 to
1·5 GPa. Therefore, it may be concluded that the effect of
H2O on the pressure estimation for the Kita-Matsuura
basalt is much less than 0·5 GPa, if the magma generation
depth ranges from 1·5 to 3 GPa.
Further information on the effect of water comes from
experiments seeking the conditions at which a primary
melt is multiply saturated with four-phase lherzolite or
three-phase harzburgite assemblages. Tatsumi et al. (1983)
experimentally showed a positive correlation between the
water content of the melt and the pressure of multiple saturation up to 2 GPa. According to their experiments,
1·5 wt % water in the melt increases the multiple saturation pressure by 0·2 GPa from 1·5 GPa under anhydrous
conditions, and 3 wt % increases the multiple saturation
pressure by 0·5 GPa from 1·8 GPa. Linear interpolation
of this relationship gives
1101
P wet ¼ P dry þ H2 Oprimary 05
:
30
ð4Þ
JOURNAL OF PETROLOGY
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JUNE 2014
Table 5: Parameters used for estimation of water content in the melt and estimated values of P,Tand H2O
Sample
Section
Group
Bulk
SiO2
Pl
Estimated values
FeO*/MgO
(wt %)
An#1
P
T
H2O2
(GPa)
(8C)
(wt %)
CFM3
H2O-CF4
(wt %)
031027c6
Kunimi
Low-SiO2
1·49
47·9
76
0·3
1173
1·02
0·32
0·70
031027e6
Kunimi
Low-SiO2
1·25
47·5
80
0·3
1193
0·83
0·17
0·69
031026a6
Kunimi
Low-SiO2
1·90
48·8
64
0·3
1140
0·88
0·32
0·60
031026a5
Kunimi
Low-SiO2
1·65
49·3
65
0·3
1191
0·23
0·34
0·15
03102147
Senryu
Medium-SiO2
1·89
51·9
72
0·3
1102
1·81
0·28
1·30
03102148
Senryu
Medium-SiO2
1·77
52·5
72
0·3
1116
1·54
0·28
1·11
03102597b
Ishimori
Medium-SiO2
1·77
50·4
73
0·3
1127
1·56
0·30
1·09
03102153
Senryu
Medium-SiO2
1·42
50·9
78
0·3
1178
0·86
0·12
0·76
03102152b
Senryu
Medium-SiO2
1·24
50·6
78
0·3
1203
0·52
0·11
0·46
03102152c
Senryu
Medium-SiO2
1·15
51·2
78
0·3
1216
0·3
0·12
0·26
031029if7
Ishimori
High-SiO2
1·04
52·6
87
0·3
1131
2·29
0·10
2·06
031029if6
Ishimori
High-SiO2
1·08
51·7
87
0·3
1143
2·2
0·13
1·92
02111276
Senryu
High-SiO2
1·36
51·5
76
0·3
1131
1·68
0·14
1·44
03101503
Senryu
High-SiO2
1·19
52·8
85
0·3
1163
1·4
0·11
1·24
03101928
Hirado
High-SiO2
1·58
51·5
76
0·3
1141
1·4
0·24
1·06
03101935
Hirado
High-SiO2
1·58
51·7
74
0·3
1158
0·98
0·24
0·74
03101924
Hirado
High-SiO2
1·33
53·7
77
0·3
1167
0·95
0·30
0·67
03101923
Hirado
High-SiO2
1·25
53·7
77
0·3
1182
0·73
0·30
0·51
03101931
Hirado
High-SiO2
1·48
51·9
72
0·3
1187
0·44
0·20
02110607
Senryu
High-SiO2
1·19
51·6
66
0·3
1229
0
0·35
0·00
*Total iron given as FeO.
1
An# [¼ 100 Ca/(Ca þ Na)] is the highest An# plagioclase core in each sample.
2
H2O content is a value estimated by assuming the bulk-rock composition as the liquid composition, which is in equilibrium with olivine with the highest Fo# and plagioclase with the highest An# in the sample.
3
CFM represents the fractionated mass ratio from the least fractionated samples of each SiO2 group.
4
H2O-CF represents the water content in the melt; values were corrected for crystal fractionation from the least differentiated samples (PM1, PM2, and PM3) for each group.
We use equation (4) to estimate the melting pressure under
hydrous conditions by comparing the estimated KitaMatsuura primary melts with the results of anhydrous
high-pressure melting experiments in a CIPW normative
nepheline^olivine^quartz (Ne’^Ol’^Qtz’) ternary projection (Irvine & Baragar, 1971), following Sakuyama et al.
(2009) (Fig. 10). The estimated melting pressure is in the
range of 2·5^2·8 GPa for the low-SiO2 group, 2^2·3 GPa
for the medium-SiO2 group, and 1·5^1·6 GPa for the highSiO2 group; these results apply if the Fo# of olivine in
equilibrium with the primary melts was 90 (Fig. 10).
Taking the effect of water into account by applying equation (3), the melting pressures of the primary magmas of
the three basalt groups are now estimated to be
2·5^2·9 GPa, 2·1^2·4 GPa, and 1·7^1·9 GPa for the low-,
medium-, and high-SiO2 groups, respectively. If we
assume a higher degree of melting, resulting in a Fo# of
91 in the residue, for the high-SiO2 group as an extreme
case, the melting pressure of the high-SiO2 group primary
melt is estimated to be 1·8^2·1GPa, which is distinctly
lower than that of the low-SiO2 group (Fig. 10). It should
be noted that the estimation of melting pressure in this
study and in that of Sakuyama et al. (2009) does not
depend on source peridotite composition, because the isobaric compositional trends of the partial melts in the Ne’^
Ol’^Qtz’ ternary do not shift significantly (Hirose &
Kushiro, 1993) even if the peridotite contains 50% of a basaltic component (Kogiso et al., 1998).
1102
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Melting temperature
There is a simple relationship between temperature, pressure, and MgO content in liquids saturated with olivine
and/or pyroxene (Ramsay et al., 1984; Sugawara, 2000;
Maaloe, 2004; Herzberg et al., 2007). We recompiled the
results of anhydrous melting experiments saturated with
olivine and pyroxene under mantle melting conditions
(0·5^4 GPa) and performed a linear regression analysis of
the dependence of the temperature (T in 8C) as a function
of both pressure (P in GPa) and the MgO content
(wt %). We obtained the following relationship for anhydrous conditions (see Appendix for details):
T ¼ 1080 7 þ 54 75 0 27074P 2
ð5Þ
þ2 21634P 0 99731 þ 15 08MgO:
The olivine liquidus temperature estimated by equation
(5) represents the maximum temperature at which the basaltic melt was generated because water in the melt decreases the olivine liquidus temperature. By combining
equations (3) and (5), we have estimated the segregation
Fig. 9. Variation diagrams for TiO2, Al2O3, and FeO* vs MgO for groundmass compositions in the Kita-Matsuura basalts (a^c) and fractional
(d^f, j^l) and equilibrium (g^i, m^o) crystallization paths modeled by the MELTS program. Calculated MELTS trends are shown for several
different initial melt water contents at 0·2 GPa, and at different crystallization pressures for fixed amounts of water (0·25, 0·50, 1 and 2 wt %
for the low-, medium-, and high-SiO2 groups) in (d^i).
(continued)
1103
JOURNAL OF PETROLOGY
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JUNE 2014
Fig. 9. (Continued).
temperature of each Kita-Matsuura basalt group from the
segregation pressure and primary melt compositions,
including the water content estimated above. Estimated
conditions are listed in Table 6 and shown in Fig. 11. The
most plausible ranges of melt segregation conditions are
summarized in Table 6.
The segregation temperatures were estimated to be
1450^15008C, 1400^14608C, and 1350^14008C for the low-,
medium-, and high-SiO2 groups, respectively. Addition of
just olivine resulted in higher temperatures and pressure
than when both olivine and clinopyroxene were added to
estimate the primary melt composition. However, even
though clinopyroxene was added with olivine, the temperature and pressure of melting clearly decrease from
low-, through medium-, to the high-SiO2 group (Fig. 11).
Therefore, the uncertainty surrounding deep cryptic fractionation of clinopyroxene does not affect the following
discussion.
Mantle potential temperature
Mantle potential temperature is one of the most important
parameters in an area of upwelling mantle (McKenzie,
1984; McKenzie & Bickle, 1988; Putirka et al., 2007; Lee
et al., 2009, and references therein). To estimate the potential temperature of the upwelling mantle from volcanic
rocks, either the source mantle composition or the degree
of melting at which the estimated primary melt was generated are prerequisites, as the adiabat of melting mantle is
different from that of solid mantle (McKenzie, 1984):
10^208C GPa1 for solid mantle and 708C GPa1 for
melting mantle (e.g. McKenzie, 1984; Iwamori et al., 1995).
The primary melt of the low-SiO2 group is most likely to
be the closest to the initial melt from the upwelling
mantle beneath Kita-Matsuura. Therefore, provided that
the primary melt of the low-SiO2 group was an initial
sample of the very first melt generated from the upwelling
mantle, we can estimate the minimum potential temperature of the upwelling mantle to be 14508C by extrapolating the adiabat of a solid mantle (Fig. 11a). We argue that
the potential temperature of the source mantle for the
Kita-Matsuura basalt was higher than 14508C, which is
1008C higher than the mantle potential temperatures
inferred beneath normal mid-ocean ridges, the Japan Sea,
and the Chugoku district in southwestern Japan
1104
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Table 6: Estimated melting conditions for the three basalt
groups
Group
Fig. 10. Normative compositions of estimated primary melts in the
ternary nepheline^olivine^quartz. Isobaric curves defining the compositions of partial melts of anhydrous peridotite are from a compilation by Sakuyama et al. (2009). The projection scheme is after Irvine
& Baragar (1971): Ne’ ¼Ne þ 0·6Ab; Qtz’ ¼Qtz þ 0·4Ab þ 0·25Opx;
Ol’ ¼Ol þ 0·75Opx.
(McKenzie, 1984; Yamashita & Fujii, 1992; Tatsumi et al.,
1994; Iwamori et al., 1995; Lee et al., 2009), 0^1008C higher
than that beneath back-arc basins (Wiens et al., 2006), and
50^2008C lower than that of a hotspot (Putirka, 2005; Lee
et al., 2009).
Temporal and spatial changes in melting
conditions
Now that we have estimated the melting conditions for the
primary melts, the temporal and spatial changes of the
magma groups shown in Fig. 1b can be interpreted as reflecting temporal and spatial variations in the melting conditions (Fig. 11b). Basaltic volcanism began at the Hirado
section in the westernmost part of the Kita-Matsuura
area. Beneath each section, the estimated melting pressure
and temperature decrease over time, whereas the water
content of the primary melt increases until eruption of the
high-SiO2 group, which is generated at the shallowest
level in the mantle. Magma eruption ceased slightly earlier
(7·5 Ma) in the Hirado section than in the other sections
(Sakuyama et al., 2009). Meanwhile, the Kunimi section,
in the easternmost part of Kita-Matsuura, 35 km east of
the Hirado section, produced only melt generated at high
pressure and temperature under virtually anhydrous conditions (low-SiO2 group) and the duration of the basaltic
volcanism was shorter than at the other sections (Fig. 11b).
Basalt generated at higher pressure and temperature
under nearly anhydrous conditions is more abundant in
the marginal sections such as Ishimori and Kunimi than
in the Hirado and Senryu sections. This spatial variation
can be extended to the west of the Hirado section, as the
basalt in the Ikitsuki section is more similar to those in sections to the east of the Hirado section than that in the
Hirado section in terms of eruption age and major element
chemical composition (Fig. 11b). Assuming that melting
Fo#
H2O
Phase*
P (GPa)
T (8C)
High-SiO2y
90
1·5
ol
1·9
1370
High-SiO2
90
0·6
ol
1·7
1398
High-SiO2
90
0·0
ol
1·7
1418
High-SiO2
90
1·5
ol þ cpx
1·7
1346
High-SiO2
90
0·6
ol þ cpx
1·6
1373
High-SiO2
90
0·0
ol þ cpx
1·5
1392
High-SiO2
91
1·5
ol
2·1
1424
High-SiO2
91
0·6
ol
2·0
1451
High-SiO2
91
1·5
ol þ cpx
1·9
1387
High-SiO2
91
0·6
ol þ cpx
1·8
1412
Medium-SiO2y
90
0·9
ol
2·4
1454
Medium-SiO2
90
0·3
ol
2·3
1460
Medium-SiO2
90
0·0
ol
2·3
1470
Medium-SiO2
90
0·9
ol þ cpx
2·2
1418
Medium-SiO2
90
0·3
ol þ cpx
2·1
1424
Medium-SiO2
90
0·0
ol þ cpx
2·0
1433
Low-SiO2y
90
0·5
ol
2·9
1483
Low-SiO2
90
0·0
ol
2·8
1499
Low-SiO2
90
0·5
ol þ cpx
2·6
1449
Low-SiO2
90
0·0
ol þ cpx
2·5
1465
*Phase indicates added crystal(s) to estimate primary melt
compositions. Ol and ol þ cpx represent olivine and olivine þ clinopyroxene addition, respectively.
yThe most plausible melting conditions. Melting conditions
assuming Fo# ¼ 91 for the high-SiO2 group are also
included to discuss the effects of the extent of source depletion. (See main text for details.)
Melting pressure and temperature were estimated by the
method described in the main text assuming Fo# of olivine
in equilibrium with the primary melt and H2O content in the
primary melt.
pressures under hydrous conditions changed from 2·9 GPa
(low-SiO2) to 1·9 GPa (high-SiO2) from 7·7 to 6·5 Ma in
the Senryu section, we propose mantle upwelling at a velocity of 2 cm a1.
DISCUSSION
Sakuyama et al. (2009) proposed that the Kita-Matsuura
basalts could be separated into low-, medium- and highSiO2 groups that exhibit separate liquid lines of descent.
By applying the melting model of Ozawa & Shimizu
(1995) to fluid-immobile elements, Sakuyama et al. (2009)
concluded that the degree of melting systematically
increased from the low-, through medium-, to high-SiO2
groups. The volcanism started at the Hirado section with
an eruption of medium-SiO2 group basalt followed by
high-SiO2 group basalt. The volcanism at the Senryu and
1105
JOURNAL OF PETROLOGY
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JUNE 2014
Fig. 11. (a) Melting temperatures and pressures estimated for primary melts with various water contents for the low- (open and filled circles),
medium- (open and filled triangles), and high-SiO2 (open and filled squares) basalt groups in equilibrium with olivine of Fo# 90. The results
for a residual olivine composition of Fo# 91 are also shown for the high-SiO2 group defined by the diagonally striped rectangle. The water content is represented by the size of the filled symbols, and dry conditions are defined by open symbols. Higher-pressure primary melts were estimated from olivine addition, whereas lower-pressure melts were estimated from olivine þ clinopyroxene addition. Experimentally determined
solidi (KLB1 and HK66; Takahashi & Kushiro, 1983; Takahashi et al., 1993) and compiled solidi for variable water contents (Hirschmann,
2000; Katz et al., 2003) are also shown. Adiabatic P^T paths for solid peridotite (178C GPa1) and melting peridotite (708C GPa1) for a
mantle potential temperature of 14508C are shown by the thick gray lines after McKenzie (1984). The horizontal lines with double arrowheads
indicate the range of potential temperatures for the Chugoku district, southwestern Japan (Iwamori, 1991) and mid-ocean ridges (MOR) (Lee
et al., 2009). The bold dotted-line arrow indicates a hypothetical melting adiabat for the low-SiO2 group primary magma. (b) Estimated melting
pressures and temperatures are shown for the five Kita-Matsuura basalt localities; bold dashed and fine dotted lines define isochrons to clarify
the spatio-temporal changes. The conditions for Takushima and Ikitsuki were estimated using the whole-rock analyses of Uto et al. (2004).
Gray areas represent interpolated melting paths.
1106
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Ishimori sections followed, with eruption of low-SiO2
group magma progressing to medium- and high-SiO2
group magma, whereas the Kunimi section produced
low-SiO2 group magma all the way up to the uppermost
horizon. Using new data, we have further constrained the
magma generation conditions beneath Kita-Matsuura: (1)
the water contents of the primary magma in the low-,
medium-, and high-SiO2 groups are estimated to be 0·5,
0·9, and 1·5 wt %, respectively; (2) the melting pressure
decreased from 2·5^2·9 to 2·1^2·4, and 1·7^1·9 GPa and
the temperature decreased from 1450^1500 to 1400^1460
and 1350^14008C for the low-, medium-, and high-SiO2
groups, respectively; (3) the estimated mantle potential
temperature was 14508C or higher, even after taking
into account the effects of water in the peridotite system.
The Kita-Matsuura basalts have the following geochemical characteristics: (1) HFSE/LREE decreases with
decreasing HFSE abundance (Fig. 4a); (2) LREE/HREE,
MREE/HREE, Zr/Y (Fig. 4c), and Nb/Y (Fig. 4d), and
Nb/Th decrease from low-, through medium-, to the
high-SiO2 group. Combining these petrological and geochemical results with previously published geological data
(Sakuyama et al., 2009), we can now constrain the melting
processes beneath northwestern Kyushu from 9·5 Ma until
c. 6 Ma.
Melting of a homogeneous mantle source
The decrease of HFSE/large ion lithophile element (LILE)
and HFSE/LREE ratios with time and with decreasing
HFSE abundance are common features in all the studied
sections in the Kita-Matsuura area. The relative depletion
of HFSE with respect to LILE and LREE has been generally attributed to (1) the presence of residual accessory
minerals (e.g. rutile, titanite, zircon) that retain HFSE in
the melting residue (e.g. Wood et al., 1979; Saunders et al.,
1980), (2) interaction of the ascending magma with surrounding harzburgitic peridotite (e.g. Kelemen et al.,
1990), or (3) melting of a mantle source composition modified by the influx of LILE-enriched slab-derived fluids or
melts (e.g. Brenan et al., 1994). Melting models that
assume a homogeneous mantle source in a system closed
to input (Fig. 12b, model 1), however, fail to reproduce the
systematic coupling of an increase in the water content in
the primary magma and the increase in the magnitude of
the HFSE depletion relative to LILE and LREE with a decrease in the trace element abundance, because water and
HFSE, LILE and REE are all incompatible elements.
Thus, either a heterogeneous source mantle (Fig. 12b,
model 2) or the involvement of an agent rich in water and
fluid-mobile components (Fig. 12b, model 3) is required.
Melting of internally heterogeneous
upwelling mantle
The source mantle may be chemically heterogeneous on a
scale much smaller than 10 km (Fig. 12b, model 2), which
is the distance between the studied sections in the KitaMatsuura area. In this model, each SiO2 group would
have formed by the melting of a specific portion of the heterogeneous mantle and segregation of the resultant melt
at an appropriate depth. However, melting and melt segregation from a certain portion of the mantle will suppress
further melting of this portion at shallower levels.
The important basic premise of this model is that the
cross-over of the solidus in upwelling mantle is coupled
with the timing of melt segregation without further melt
generation. The onset of melting of a hydrous mantle peridotite source that rises adiabatically in a closed system is
deeper than that for the equivalent anhydrous mantle because the solidus temperature is lowered in the presence of
water. Therefore, melt production in the deeper portion of
the upwelling mantle should be higher for hydrous
peridotite.
Because the degree of melting increases from the low- to
the high-SiO2 group (Sakuyama et al., 2009), as well as
the water content in the primary magma, the water content in the source mantle should also increase from the
low- to the high-SiO2 group (Hirose & Kawamoto, 1995;
Katz et al., 2003). It is thus difficult to explain why the
high-SiO2 group magma with a high H2O content was
derived from the shallowest depth. Although a refractory
mantle peridotite composition can delay the onset of melting, the estimated highest degree of melting of the highSiO2 group of the Kita-Matsuura basalts cannot be
achieved by this model. This is because the degree of melting at a shallow level is more strongly controlled by the
source depletion rather than its H2O content. This is due
to the decreased role of H2O in lowering the solidus temperature of a peridotite at a greater degree of melting by
the dilution of H2O in the partial melt.
This suggests that water must have been introduced into
the source mantle of the high-SiO2 group after the segregation of the medium-SiO2 melt, violating the assumption
of a closed system. We therefore conclude that the essential
features of the Kita-Matsuura basalts cannot be produced
by any internally heterogeneous mantle model, but instead
that they require a supply of aqueous fluid into the melting
system at each melting stage.
Progressive melting of a homogeneous
upwelling mantle source and addition of
fluid from a chemically stratified mantle
Another plausible melting model is the progressive addition of H2O-bearing fluid to an upwelling of mantle
source (Fig. 12b, model 3). This is the most plausible
model that can be used to explain the progressive increases
in H2O and LILE contents, and the degree of melting
with simultaneous decreases in pressure and temperature
over time. There are multiple mechanisms to achieve the
process described above. The first model involves a
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Fig. 12. (a) Schematic illustration of the diapiric upwelling model for the petrogenesis of the Kita-Matsuura basalts. (b) Schematic illustrations
of the melting models: model 1, progressive decompression melting of a single source mantle in a system closed to input; model 2, decompression
melting of an internally heterogeneous mantle, each portion of which provides melt of low-, medium-, and high-SiO2 basalt composition independently (see the main text for details); model 3, progressive decompression melting open to both input and output. (c) Schematic illustration
of a homogeneous diapiric upwelling mantle that is interacting with a geochemically distinct hydrous ambient mantle for each instant of time
shown in (b).
progressive supply of fluids from the downgoing Philippine
Sea slab to adiabatically upwelling anhydrous mantle, as
occurs in the mantle wedge above a subduction zone. The
second model involves the interaction of an ascending
mantle diapir with a geochemically distinct shallower
upper mantle layer as proposed by Sakuyama et al. (2009).
Although recent seismic topography studies have revealed
the presence of a downgoing slab beneath northwestern
Kyushu (Zhao et al., 2012), as is the case with our previous
model, we prefer the second model. This is because the
imaged high Vp anomaly beneath northwestern Kyushu is
not as clear as those imaged beneath other subduction
zones, and the chemical compositions of Cenozoic basalts
in northwestern Kyushu are clearly distinct from those of
typical subduction zone volcanism. Upwelling of the
mantle in this region is suggested to be related to the initiation of the opening of the Beppu-Shimabara Rift at the
northern extension of the Okinawa Trough, which
occurred before 6 Ma. The presence of pre-metasomatized
mantle beneath north Kyushu is also reasonable because
the area had long been a supra-subduction zone setting,
related to the subduction of the Pacific Plate slab since the
Cretaceous up to the Middle Miocene before opening of
the Japan Sea (60^15 Ma; e.g. Kimura et al., 2005).
When part of the cooler ambient mantle is mechanically
incorporated into the thermally buoyant mantle diapir, it
will be heated and decompress concurrently as it rises
within the diapir, forming a mushroom-like morphology
(Griffiths, 1986). If the incorporated mantle is heated to
temperatures high enough to melt or dehydrate, it will provide a melt or fluid to the overlying original diapiric
mantle (Fig. 12b, model 3 and Fig. 12c). This process may
promote further melting of the original upwelling diapir,
even if the upwelling mantle becomes refractory after
melt segregation. Below, we discuss a mantle melting
model open to both outputs and inputs to explain the petrological, geochemical, and spatiotemporal variations in
the Kita-Matsuura basalts.
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SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Multistage open-system melting model
To account for the addition of fluid into the melting system,
we used the open-system melting model of Ozawa (2001),
which can model simultaneous influx of melt or fluid and
melt segregation in multiple stages. The equation used is
from equation (59) of Ozawa (2001):
for f5fc ¼ ac/[1 þ b(1 ac)],
lq,in
lq
Cj
¼
Cj0 þ Cj
bf
Pj0 þ f ð1 þ b Pj Þ
ð6Þ
and for f fc ¼ ac/[1 þ b(1 ac)],
h
i9
8
lq,in,k
lq,in,k
lq,0,k
>
b k Cj
bk Cj
Cj
ð1 Pjk þ bk Þ >
>
>
>
>
>
>
<
=
k þbk
1
1P
lq
j
"
#
Cj ¼
k
k
k
k
>
f k ð1 Pjk þ b gk Þ 1Pj þb g
1 Pjk þ bk >
>
>
>
>
>
>
: 1 þ 0,k
;
0,k
0,k
Pj ð1 a Þ þ a
ð7Þ
lq
Cj
is the concentration of the jth component in
where
lq,in
is the concentration of the jth component
the melt, Cj
lq,0
in the influxing fluid, Cj is the initial concentration of
the jth component in the melt, a is the melt fraction in the
initial system, ac is the critical melt fraction, b is the mass
influx rate (influxing mass fraction of the initial solid
mass divided by the degree of melting), f is the extent
of solid^melt reaction relative to the initial solid mass,
Pj is the weighted average of the jth partition coefficient
between the solid and melt, P j0 is the initial bulk concentration of the jth component normalized to the initial
melt composition, and g is the mass separation rate
(separated mass fraction of the initial solid mass divided
by the degree of melting). Parameters with superscript k
are the parameters for the kth stage. In this model, influx
begins simultaneously with the initiation of melting. A
schematic illustration of the key parameters is shown
in Fig. 13.
In our previous model, only single-stage critical melting
was examined for Nb/Yand Zr/Y (Sakuyama et al., 2009).
However, the necessity of a melting model also open to
input is required to explain the temporal changes in pressure, temperature, and water content in the melt consistently (see above). The open-system critical melting
(OSCM) model of Ozawa & Shimizu (1995), which assumes constant influx and melt separation rates during
the entire melting event, can handle more general opensystem problems than the model of Sobolev & Shimizu
(1993). The assumption of constant influx and melt separation rate is, however, inappropriate for the KitaMatsuura basalts, particularly because the HFSE/LILE
ratio of the samples decreases with decreasing HFSE abundance, together with a temporal increase in water content
in the estimated primary melts (Table 6). This strongly suggests increases of water-rich fluid or melt influx into the
Fig. 13. Schematic illustration of the open-system melting model used
in this study. (See text for details.)
melting system as melting proceeds. We therefore employed
the multistage melting model represented by equation (7),
in which the extent to which the system is open can be
varied (by adjusting b and g), although the model requires
more parameters (e.g. number of stages) that are difficult
to constrain. However, this choice as a conceptual basis
should be more reasonable compared with other models.
We separated the melting process into six stages (Table 7
and Fig. 14a). The first three stages involve garnet peridotite melting, the fourth stage melting during the transition
from garnet to spinel peridotite, and the last two stages
spinel peridotite melting. The first stage allows only input
into the source to investigate the degree of enrichment of
the original source before melting. Therefore, we set up
two stages for both the garnet and spinel stages and one
stage for the transition from garnet to spinel. Because melting started at the garnet stage and finished at the spinel
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Table 7: Estimated set of parameters for open-system melting at each melting stage and assumed mineral and melting mode
Stage
Enriched mantle source (Workman et al., 2004)
Initial value at each melting stage
f
Lithology
Mode (wt fraction)
Ol
Opx
Cpx
Sp
Parameters for open-system melting
a
b
g
Input
Output
Gt
1
0·000
0·550
0·200
0·150
0
0·100
Gt lhz
0·000
0·020
0·00
Closed
Closed
2
0·010
0·555
0·204
0·143
0
0·098
Gt lhz
0·010
0·020
0·50
Open
Open
3
0·025
0·562
0·210
0·133
0
0·095
Gt lhz
0·018
0·040
0·90
Open
Open
4
0·045
0·572
0·218
0·119
0
0·091
Gt/Sp lhz
0·021
0·070
1·10
Open
Open
5
0·063
0·564
0·296
0·159
0·021
0
Sp lhz
0·021
0·100
1·10
Open
Open
6
0·100
0·599
0·292
0·132
0·019
0
Sp hzb
0·020
0·140
1·20
Open
Open
Melting mode
Ol
Gt stage
Sp–Gt transition
Sp stage
Opx
Cpx
Sp
Gt
0
0·3
0·08
–0·19
0·81
1·00
–1·90
–3·70
–1·07
–0·30
0·40
0·82
0·08
4·67
0
Range of melting degree and melting stage for accumulated melt to reproduce each group
Z
Degree
Stage
Lithology
Low
0–0·065
1–5
Gt-lhz–Gt/Sp-lhz
0·14 (wt %)
Medium
0·025–0·090
3–5
Gt-lhz–Sp-lhz
0·27 (wt %)
High
0·045–0·120
4–6
Sp-lhz–Sp-hzb
0·49 (wt %)
Primitive mantle source (Sun & McDonough, 1989)
Stage
Initial value at each melting stage
f
Lithology
Mode (wt fraction)
Ol
Opx
Cpx
Sp
Parameters for open system melting
a
b
g
Input
Output
Gt
1
0·000
0·550
0·200
0·150
0
0·100
Gt lhz
0·000
0·010
0·00
Closed
Closed
2
0·015
0·557
0·206
0·140
0
0·097
Gt lhz
0·015
0·025
0·50
Open
Open
3
0·020
0·560
0·208
0·137
0
0·096
Gt lhz
0·018
0·035
0·80
Open
Open
4
0·035
0·567
0·214
0·126
0
0·093
Gt/Sp lhz
0·021
0·050
1·10
Open
Open
5
0·054
0·558
0·293
0·167
0·022
0
Sp lhz
0·021
0·080
1·10
Open
Open
6
0·100
0·602
0·288
0·134
0·019
0
Sp hzb
0·021
0·090
1·10
Open
Open
Melting mode
Ol
Gt stage
Sp–Gt transition
Sp stage
Opx
Cpx
Sp
0·08
–0·19
0·81
1·00
–1·90
–3·70
–1·07
0
–0·30
0·40
0·82
0·08
Gt
0·3
4·67
0
Range of melting degree and melting stage for accumulated melt to reproduce each group
Z
Degree
Stage
Lithology
Low
0–0·041
1–4
Gt-lhz–Gt/Sp-lhz
0·10 (wt %)
Medium
0·020–0·058
3–5
Gt-lhz–Sp-lhz
0·20 (wt %)
High
0·035–0·185
4–6
Sp-lhz–Sp-hzb
1·31 (wt %)
f, a, b, g, and Z are the degree of melting, the mass fraction of melt in the current system of the melting stage, the influx rate, the
separation rate, and the ratio of total influxed material relative to the initial solid mass, respectively. Ol, Opx, Cpx, Sp, Gt, lhz, and hzb
are olivine, orthopyroxene, clinopyroxene, spinel, garnet, lherzolite, and harzburgite, respectively. Low, medium, and high represent low-,
medium-, and high-SiO2 groups, respectively. Estimated Z is the value at the highest degree of melting for each magma group.
1110
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Fig. 14. Results of the open-system melting model. (a) Optimum model parameters for the mineral mode, melt influx rate (b) and total influxed mass relative to the initial mass as a function of the degree of melting. The range of accumulated melt in the primary melts of each
basalt group and the degree of melting at which the garnet^spinel transition occurs are also indicated. (b) Primitive mantle normalized trace
element patterns of altered oceanic crust (AOC; gray field), sediments (between continuous lines), average of Pacific Plate sediments (filled diamonds), and average of Pacific Plate AOC (open diamonds). Also shown is the calculated fluid composition derived from AOC. Modelled
primitive mantle normalized trace element patterns for the primary melts of (c) high-SiO2, (d) medium-SiO2, and (e) low-SiO2 basalt
groups from a PMTL source (filled diamonds) and an EM source (open diamonds). In these panels the gray fields define the trace element patterns of the estimated primary melts. The instantaneous melt composition from PMTL source mantle is shown by a bold dotted line. Each primary melt targeted for use in the model calculations is shown by a dashed line.
stage, we had to take into account the effect of the phase
transition during modelling. Although the number of
stages for the multistage melting model is arbitrary, this
configuration is rather simple to treat melting from the
garnet to the spinel stage comprehensively. Mineral
modes and melting parameters such as a, b, and g can be
different for each stage, but are assumed to be constant
within each stage. The initial mineral mode and the melting modes for the garnet and spinel peridotite melting
stages are listed in Table 7, and the initial composition is
assumed to be primitive mantle (PMTL; Sun &
McDonough, 1989) and enriched mantle (EM; Workman
et al., 2004), the compositions of which are given in
Table 8. Sakuyama et al. (2009) have already shown that a
depleted MORB-source mantle (DMM) cannot reproduce
the HFSE compositions of the Kita-Matsuura basalts. The
melting stoichiometries for the garnet and spinel peridotite
melting are the same as those adopted by Sakuyama et al.
(2009) and the initial bulk mineral mode follows Johnson
et al. (1990). Reaction stoichiometry for the transition from
garnet to spinel peridotite is according to Walter et al.
(1995). We ignored any temperature and pressure dependence of partition coefficients between mineral and fluid or
melt, the effect of which will be discussed below.
We explored other melting parameters (e.g. F, a, b, and
g) that could reproduce the estimated primary melt compositions by accumulated melting. First, we estimated an
optimum parameter set at the degree of melting estimated
for the low-SiO2 group (lowest degrees of melting). Then,
we estimated the sets of parameters for successive melting
stages corresponding to the medium- and high-SiO2
groups (at higher melting degrees) starting from the residue of the previous melting stage. This treatment is based
on our model concept that assumes progressive melting of
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Table 8: Partition coefficients of minerals and chemical compositions of initial mantle and influxed material assumed in
open-system melting calculations
Ol
Opx
Cpx
Sp
Gt
PMTL
EM2
Rb
0·000179
0·0006
0·00175
0
0·0007
0·635
1·456
261·4
Ba
0·000032
0·0035
0·0006
0
0·0007
6·989
15·857
2051·8
Th
0·000052
0·013
0·00531
0
0·001
0·085
0·177
22·7
U
0·000018
0·0017
0·00361
0
0·00918
0·02
0·040
4·5
Ta
0·00007
0·003
0·0102
0·02
0·02
0·041
0·063
0·6
Nb
0·0001
0·003
0·0077
0·02
0·015
0·71
1·087
17·2
K
0·000177
0·0003
0·0072
0
0·0007
La
0·000028
0·0025
0·0536
0·01
0·0015
0·687
0·895
Ce
0·000038
0·005
0·0858
0·01
0·008
1·775
1·923
62·8
Pb
0·000479
0·0013
0·072
0
0·0005
0·185
0·144
109·3
Pr
0·0008
0·0048
0·15
0·01
0·054
Sr
0·0015
0·007
0·1283
0
0·006
Nd
0·00042
0·0068
0·1873
0·01
0·087
Zr
0·007
0·021
0·123
0·02
0·5
Hf
0·0038
0·01
0·256
0·02
0·24
Sm
0·0013
0·01
0·291
0·008
0·7
Eu
0·0016
0·013
0·31
0·007
Gd
0·0055
0·016
0·3
Tb
0·0041
0·019
Dy
0·01
Ti
Y
250
Influx
0·276
21
1·354
50431·5
33·8
0·251
2·5
20·044
253·0
1·140
23·4
8·835
109·4
0·31
0·238
2·2
0·444
0·347
4·9
0·9
0·17
0·128
1·6
0·006
1·19
0·6
0·445
4·9
0·31
0·009
1·5
0·108
0·082
0·7
0·022
0·33
0·01
2·2
0·74
0·569
5·2
0·006
0·024
0·384
0·048
0·65
0·007
0·06
0·421
0·0023
2·8
4·55
3·655
28·2
Ho
0·007
0·026
0·31
0·009
3·3
0·164
0·127
1·0
Er
0·0087
0·03
0·29
0·01
3·6
0·48
0·378
3·2
Tm
0·009
0·12
0·255
0·01
3·5
0·074
11·2
1300
900·5
5427·5
Yb
0·017
0·049
0·28
0·008
3·88
0·49
0·387
3·2
Lu
0·02
0·06
0·28
0·02
3·79
0·07
0·061
0·5
Ol, Opx, Cpx, Sp, Gt, PMTL, and influx represent olivine, orthopyroxene, clinopyroxene, spinel, garnet, primitive mantle (Sun
& McDonough, 1989), and assumed influx fluid, respectively. EM2 is from Workman et al. (2004). Values in the columns
for minerals are the assumed partition coefficients, and those for PMTL, EM2, and influx are concentrations in ppm.
a common source mantle that consecutively produces the
three basalt groups as the mantle ascends.
Composition of the influxing material
As shown by Sakuyama et al. (2009) and from the newly
obtained geochemical data for the Kita-Matsuura basalts,
LILE and LREE enrichment relative to HFSE is more evident for the medium- and high-SiO2 basalt groups, which
suggests that the contribution of a fluid released from the
subducted slab should be considered along with water
added to the source mantle beneath Kita-Matsuura.
Because the upper mantle beneath northwestern Kyushu
was located above the subduction zone before the opening
of the Japan Sea (Kimura et al., 2005), the upper mantle
must have been affected by fluid released from the subducting slab.
When subducting slabs dehydrate and release aqueous
fluid, large fractions of the LILE, LREE, Th, U, and Pb
are efficiently removed from the subducted sediments and
altered oceanic crust. The chemical composition of the
released fluid depends on the source material and mobility
of each element:
mobility ¼
CSTM CRP
100ð%Þ
CSTM
ð8Þ
where CSTM and CRP are an element concentration of the
starting material and run products, respectively (Tatsumi
et al., 1986).
Primitive mantle normalized trace element patterns of
fluids from subducting sediments generally show a relatively flat pattern from Rb to U, with strong positive Pb
and K anomalies, and a weak positive Ti anomaly (Plank
1112
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
& Langmuir, 1998; Plank et al., 2007; Fig. 14b) similar to
the patterns of the Kita-Matsuura basalts, whereas fluids
derived from dehydration of altered oceanic crust show
strong depletion in Ba, Th, and Pb (Alt & Teagle, 2003;
Kelley et al., 2003; Nakamura et al., 2007; Fig. 14b). Thus
fluids derived from subducting sediment may dominate
the upper mantle signature beneath the back-arc region of
Kyushu.
There are two candidates for the origin of the sedimentderived fluids, one from the Pacific Plate slab and the
other from the Philippine Sea Plate slab. Although the
trace element compositions of subducted sediments are different between the two oceanic plates (e.g. Plank &
Langmuir, 1998), the difference is much smaller than the
difference between sediment and altered oceanic crust.
Thus, it is difficult to distinguish the origin of the fluid
without additional isotopic data. Therefore, first, we
assumed that sediment with an average composition of
sediments from the Pacific Plate and Philippine Sea Plate
[reported by Plank & Langmuir (1998)] was the hypothetical sediment that releases fluid.
As a first-order approximation, we used a slab sediment
fluid composition derived by 1% dehydration of Pacific
Plate sediment in accordance with the mobilities of trace
elements determined by Aizawa et al. (1999). There are a
number of ways to transport the sediment-derived fluid
component into the upwelling anhydrous mantle. These include (1) direct addition of the fluid through channel flow
from the Pacific Plate slab to the upwelling mantle and (2)
entrainment of fluid-metasomatized, water-bearing
mantle (either lithospheric or asthenospheric) in the upwelling dry mantle. The latter step needs an additional
process that includes sediment fluid metasomatism, fluid
release from the metasomatized mantle, and introduction
of the secondary fluid into the dry mantle. It is impossible
to treat such a complex process quantitatively. We therefore
used the immediate sediment fluid composition and tested
whether or not such a fluid can account for the opensystem melting model. The calculated chemical composition of this influxing material into upwelling mantle is
given in Table 8. Below we discuss the calculation results
for open-system melting. The effects of the composition of
the influxed material on the estimation of b are also discussed below.
Modelling results
The parameters used in the modelling are listed in Table 8
and the calculated trace element compositions are shown
in Fig. 14. Sakuyama et al. (2009) proposed that the lowSiO2 group can be reproduced by the lowest degree of
melting whereas the high-SiO2 group can be reproduced
by the highest degree of melting. The calculated degree of
melting generally agrees in order, but with slightly different values (Table 8 and Fig. 14). The origin of this
difference is related to the difference of the adopted melting models, which is discussed further below.
Optimized trace element compositions of the partial
melt from both the PMTL and EM source reproduce the
primary melt for each group reasonably well. In particular, the LILE- and LREE-enriched patterns of the lowSiO2 group and enriched patterns with depletions in
HFSE for the medium- and high-SiO2 groups are reproduced by both PMTL and EM source compositions.
Differences between these source materials can be partially
compensated for by different degrees of melting and the
amount of influxed material. The PMTL source gives relatively better results (Fig. 14c^e). For the PMTL source the
degrees of melting for the low- and medium-SiO2 groups
need to be lower and that for the high-SiO2 group higher
than those for the EM source in order for the PMTL
source to produce a melt more enriched in LILE and
LREE and depleted in HREE than the EM source at the
garnet stage. A higher degree of melting is required for
the PMTL source to produce a melt with more depleted
HREE at the spinel stage than for the EM source.
Melting of the PMTL source requires more fluid influx
than the EM source because LILE and LREE are more
enriched in the EM source than in the PMTL source. The
EM source gives a lower MREE abundance for the lowSiO2 group, and lower LREE and MREE for the
medium-SiO2 group. Both the PMTL and EM sources
yielded slightly higher HREE abundances for the
high-SiO2 groups because the contribution of the melt generated at the spinel stage is greater at higher degrees of
melting. The abundance of Pb for every SiO2 group
(Fig. 14c^e), which strongly depends on the assumed composition of the influx material, is not reproduced well by
either PMTL or EM melting.
We will discuss the results for the PMTL source in more
detail below. To reproduce the trace element patterns of
the low-SiO2 group primary melt, a minor influx of material is required (b ¼0·010^0·035) during the garnet peridotite melting stages, up to 4·5% melting, and a further
small influx of material (b ¼0·05) is necessary when melting occurs during the transition from garnet to spinel peridotite (Table 7a). Initially, once spinel peridotite is stable
(stage 5), a moderate amount of material influx (b ¼0·08)
is required during melting to reproduce the trace element
pattern of the medium-SiO2 group primary melt at
f ¼ 5·8% (Table 7a). Later, during spinel peridotite melting (stage 6), a higher influx of material (b ¼0·09) is
required to reproduce the trace element pattern of the
high-SiO2 group primary melt at f ¼ 18·5%. The optimized influx rate (b) increases from 0·01 to 0·09 between
melting to generate the low-SiO2 group and the highSiO2 group primary melts (Table 7a), which suggests that
the fluid or melt influx steadily increased as the diapir ascended. The total amount of material that was added to
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the diapir for each stage of the melting was 0·10 wt % for
the low-, 0·20 wt % for the medium-, and 1·3 wt % for the
high-SiO2 groups relative to the initial mass of the source
peridotite, which was 1·4 wt % for the low-, 4·8 wt % for
the medium-, and 7·6 wt % for the high-SiO2 group relative to the accumulated melt mass. The increases in the
rate of influx and the mass of material that influxed into
the system during the generation of the low-, through the
medium-, to the high-SiO2 primary melts are consistent
with the increase in the water content of the primary
melts estimated above, if the influxing material contained
tens of weight per cent water.
Identification of isotopic components
involved in melting and their relative
contribution
The Pb isotope compositions of volcanic rocks, especially
for subduction-related volcanism, have been used as a
tracer to investigate the contribution of fluid in the
magma generation process (e.g. Nakamura et al., 2008;
Straub et al., 2009), as Pb is considered to be highly partitioned into the fluid phase released from the igneous and
sedimentary layers of the subducting slab (Kogiso et al.,
1997; Aizawa et al., 1999; Kessel et al., 2005). The observed
systematic relationships among melting pressure, trace
element composition, and Pb isotope composition for the
Kita-Matsuura basalts allow us to estimate the contribution of appropriate source materials identified as endmember components in 206Pb/204Pb^207Pb/204Pb^208Pb/
204
Pb isotopic space.
Because the Pb isotope compositions of the KitaMatsuura basalts are highly variable (Fig. 7c^f), at least
three isotopic end-member components are required. As
mentioned above, the low-SiO2 group lavas that erupted
in the earliest stage in each section were produced by the
smallest degree of melting and were the least affected by
the influx of fluid, which allows us to assume that the Pb
isotope composition of these lavas should not have been
significantly modified from the composition of the original
upwelling mantle. However, even within the low-SiO2
group, there are systematic differences in Pb isotope composition. Therefore, we used the average composition of
the three low-SiO2 group samples with the highest
208
Pb/204Pb (Kunimi, 031026a8; Ishimori, 03102475;
Senryu, 03102145) among the Kita-Matsuura basalts as the
initial composition of the upwelling mantle: component
C1 (Fig. 15).
A sample with the lowest 208Pb/204Pb (Kunimi,
031027e4) was also classified as belonging to the low-SiO2
group, even though it has a considerably different Pb isotope composition from those of the low-SiO2 group lavas
in the earliest stage (component C1). This suggests that
the original upwelling mantle may have already been isotopically heterogeneous before fluid influx. Extension of
NUMBER 6
JUNE 2014
the trend from component C1 to 031027e4 points towards
the DMM composition (Fig. 7b, d, and f), which is inferred
to be the dominant asthenospheric mantle component.
Thus, we assume that DMM is another end-member component for the Kita-Matsuura basalts: component C2
(Fig. 15). The original upwelling mantle before it was affected by fluid is, therefore, assumed to be a mixture of
components C1 and C2. Samples of the medium-SiO2 and
high-SiO2 groups that erupted in the middle stage
(03101930, 03101924, 03102155, 02111173B, 02111276, and
03102481) are higher in 143Nd/144Nd and lower in 87Sr/86Sr
and 208Pb/204Pb than those of component C1. This suggests
that the contribution of C2 temporally increased.
Because the source of the high-SiO2 basalt is estimated
to have been affected by fluid from the subducted slab, as
discussed above, the third component should represent
the fluid influxed to the upwelling mantle. High-SiO2 samples that erupted in the latest stage of the Hirado, Senryu,
and Ishimori sections (03101817, 02110934, 02110723, and
031029if7) are lower in 143Nd/144Nd and higher in
87
Sr/86Sr, 206Pb/204Pb, and 208Pb/204Pb than the samples
erupted in the middle stage. This suggests that the effects
from the third component, which was richer in radiogenic
Pb and Sr and poorer in radiogenic Nd, became much
stronger in the latest stage of the Kita-Matsuura volcanism. The Pb isotope compositions of altered oceanic crust
and sediment on the Philippine Sea Plate and Pacific
Plate are shown in Figs 7 and 15. Subducting materials to
represent the Philippine Sea Plate were chosen from sites
in its northern part [sites 442, 443, and 444 of Deep Sea
Drilling Project (DSDP) Leg 58 and site 582 of DSDP
Leg 87]. The sediments are enriched in radiogenic Pb and
Sr and depleted in radiogenic Nd, which is consistent
with the characteristics of the third component required
in the petrogenesis of the Kita-Matsuura basalts. Altered
oceanic crust and terrigenous sediment of the Philippine
Sea Plate have higher 8/4 and 7/4 than those of the
Pacific Plate (Hickey-Vargas, 1991; Plank & Langmuir,
1998; Shimoda et al., 1998; Hauff et al., 2003; Plank et al.,
2007). Here, we assumed an average of the high-SiO2
group samples with the highest 206Pb/204Pb as a hypothetical product produced by mixing between a partial melt
of the original upwelling mantle with C1 þC2 composition
and an influxed fluid. We refer to the fluid as component
C3 (Fig. 15a).
In the open-system modelling, the degree of melting is
mostly constrained by fluid-immobile elements such as the
HFSE and HREE, and the amount of influxed material
by fluid-mobile elements such as Rb, Ba, Th, U, K and
the LREE. Good constraints are not provided by Pb and
Sr because the assumed abundance of these elements in
the influxed material may not be appropriate.
Consequently, we estimate the concentrations of Pb, Sr,
and Nd and the isotope composition of the influxed fluid
1114
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Fig. 15. (a) 208Pb/204Pb vs 206Pb/204Pb for the Kita-Matsuura basalts and illustration of the method used to estimate the composition of the influxed fluid. The open star represents the estimated composition of the influxed fluid. The other symbols are the same as in Fig. 2d. (b)
208
Pb/204Pb vs 206Pb/204Pb, (c) 143Nd/144Nd vs 87Sr/86Sr, and (d) 143Nd/144Nd vs 206Pb/204Pb for the Kita-Matsuura basalts shown by filled diamonds. The compositional ranges for sediments (SED) and altered oceanic crust (AOC) from the Pacific slab (PAC) and the Philippine Sea
slab (PHS) are shown by dark gray areas (Hickey-Vargas, 1991; Plank & Langmuir, 1998; Shimoda et al., 1998; Hauff et al., 2003; Plank et al.,
2007). The altered oceanic crust and sediment compositions of the Philippine Sea Plate and Pacific Plate used in the mixing calculations are
indicated by symbols: crosses, plus signs, open triangles, and squares. Mixing trajectories marked off every 25 wt % between the materials of
the Philippine Sea Plate are shown in (d) as continuous curves and that for the Pacific Plate as dashed curves.
by using the degree of melting and influxed mass estimated
by trace element modelling. First, the partial melt composition from the C1 mantle was calculated by accumulating
the instantaneous melts over a range of degrees of melting
from 0·045 to 0·12 for the enriched mantle source and
from 0·035 to 0·185 for the primitive mantle without any
slab input. Trace element (Sr, Nd, and Pb) concentrations
in the initial solid composition of component C1 were
assumed as those of enriched mantle (Workman et al.,
2004) and primitive mantle (Sun & McDonough, 1989).
Next, we estimated the fluid composition by assuming
that the average value of the high-SiO2 group samples
with the highest 206Pb/204Pb represents a mixture between
the melt estimated in the first step and fluid. By applying
the mass of influxed material relative to the initial solid
mass [0·35 wt % (¼ 0·49^0·14 wt %) for the enriched
mantle and 1·21wt % (¼ 1·31^0·10 wt %) for the primitive
mantle], which were estimated by trace element modelling
and conducting mass-balance calculations (Table 7), we
estimated the trace element concentrations and the Sr^
Nd^Pb isotopic composition of the fluid derived from component C3 (Table 9).
If we adopt these three mantle components (C1, C2, and
C3), all the other samples of the Kita-Matsuura basalts
can be reproduced by mixing (Fig. 15a).
The estimated Pb isotope composition of the fluid (component C3) for both the enriched and primitive mantle
sources plots close to a mixing line between altered oceanic
crust and sediment on the Philippine Sea Plate with a
strong affinity to the sediment composition (Fig. 15). This
suggests that the fluid was originally derived from the sediment layer of the Philippine Sea Plate and that it was
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Table 9: Assumed parameters and estimated fluid isotopic composition
Mass (initial solid as 1)
Sr (ppm)
Sr/86Sr
Melt from initial
Mixed melt
solid PMTL
03101930*
0·15y
108y
0·162z
87
0·704519
7·1y
143
0·512721
0·512734
Pb (ppm)
0·3y
3·1
Pb/204Pb
0·012y
214
Nd (ppm)
Nd/144Nd
Influxed fluid
1528z
0·704482
0·704450z
13·0
86·1z
0·512747z
23·8z
Melt from initial
Mixed melt
solid EM
03101930*
0·075
99·8y
0·704519
6·7y
0·0035
214
0·704482
13·0
0·512721
0·512734
0·57y
2·1
Influxed fluid
0·0785
2661
0·704452
147·1
0·512747
34·8
206
18·2522
18·3887
18·4137z
18·2522
18·3887
18·4374
207
15·5447
15·6195
15·6332z
15·5447
15·6195
15·6461
208
38·8606
38·8063
38·7964z
38·8606
38·8063
38·7870
Pb/204Pb
Pb/204Pb
*Concentrations of Sr, Nd, and Pb were corrected to estimated values of the primary melt composition in equilibrium with
residual olivine with Fo# ¼ 90, assuming the olivine maximum fractionation model. Sr–Nd–Pb isotopic compositions of
initial solid for both the PMTL and EM sources were assumed to be average compositions of the three low-SiO2 group
samples (Fig. 15a).
yValues estimated by applying the open-system melting model in this study. Compositions for elements in the melt
produced from the initial solid were calculated by the accumulated melting model without input.
zValues estimated in this study.
involved in the generation of the medium- and high-SiO2
group basalts in the later stages of the Kita-Matsuura
basalt activity, provided that the subducted sediment composition of the Philippine Sea Plate was not much different
from the current sediment on the Philippine Sea Plate.
In addition, we estimate the relative contribution of C2
and C3 to C1 during the generation of the Kita-Matsuura
basalts (Table 10) based on the composition of C3. We can
constrain only the relative contribution of the endmember components, as the absolute value of the mixing
ratio depends on the assumed composition of the endmember components. The fraction of C2 relative to C1
temporally increased in the Kunimi, Ishimori, and
Senryu sections, except during the latest activity
(02110723) in the Senryu section; however, proportions
were almost constant in the Hirado section. The contribution of sediment-derived fluid (C3) temporally increased
in every section, except for the Kunimi section. As the
SiO2 content in the primary melt increased and the estimated melting pressure deceased, the contribution of the
sediment-derived (C3) component increased. The average
contribution of the C3 fluid to the petrogenesis of the
low-, medium-, and high-SiO2 group lavas was 0·15, 0·41,
and 1·05 wt %, respectively. These values are reasonably
consistent with the total influxed fluid mass estimated
above on the basis of incompatible trace elements.
Comparison with the previous model and
implications for polybaric melting
The degree of melting estimated in this study (4·1^18% for
PMTL) is systematically higher than that obtained by
Sakuyama et al. (2009) (3^9%). The discrepancy is due to
the melt segregation process; accumulated melts were
adopted in this study, whereas Sakuyama et al. (2009) used
instantaneous melts. An accumulated melt derived by high
degrees of fractional melting has incompatible trace element abundances similar to those of an instantaneous melt
formed by lower degrees of melting (Fig. 14c^e). Although
it is difficult to constrain the melt segregation process just
from trace elements, the results of this study may be more
plausible in that they are consistent with the major element
chemistry. An increase in the degree of melting is a linear
function of the H2O content in the partial melt (Hirose &
Kawamoto, 1995). The presence of 1wt % H2O in the
melt, which corresponds to the estimated H2O content in
the primary melt of the high-SiO2 group, increases the
extent of melting at 1GPa by 20% (Hirose &
Kawamoto, 1995) in comparison with anhydrous conditions. Provided that this relationship can also be applied at
2 GPa, the degree of melting for the high-SiO2 group
should be as high as 20%, which is consistent with the
degree of melting estimated in this study. Furthermore, the
melting model that Sakuyama et al. (2009) adopted fails to
reproduce the systematic variation of both fluid-immobile
and fluid-mobile elements. We therefore conclude that the
results of this study are more plausible than those of
Sakuyama et al. (2009). However, it should be noted that
this new melting model does not invalidate the essential
conclusions of Sakuyama et al. (2009), which show that the
melting degree increased from the low- to the high-SiO2
group, as the results of the previous study are valid as long
as only fluid-immobile HFSE are considered.
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Table 10: Estimated amount of fluid component for samples
in each section
Section
Sample
and group
K–Ar age*
C2/(C1 þ C2)
Fluid
(Ma)
(fraction)
component
(wt %)
Hirado
High
03101930
7·29
0·7
1·8
High
03101924
8·1
0·65
0·65
Medium
03101817
8·84
0·84
0·70
High
02110723
6·37
0·22
0·85
High
02110934y
6·48
–
–
Medium
02111276
6·54
0·72
0·55
Medium
02111173B
6·56
0·63
0·25
Medium
03102155
6·59
0·34
0·15
Low
03102145
7·68
0·11
0·05
High
031029if7
5·42
0·50
1·45
High
03102481
7·33
0·57
0·50
Low
03102475
7·74
0·17
0·0
Low
031027e4y
7·11
0·66
0·0
Low
031027c9x
7·26
0·89
0·55
Low
031026a8
7·76
0·01
0·15
Senryu
Ishimori
polybaric mantle melting (e.g. Hirose & Kushiro, 1998)
suggest that the MgO content of an accumulated partial
melt is almost constant for variable melting degrees
(6·4^21%). There is no significant variation in MgO content among the presumed primary melt compositions of
the three Kita-Matsuura basalt groups. Even in the case
that the Fo# of olivine in the residual peridotite in equilibrium with the high-SiO2 primary melt was 91, the MgO
content of the primary melt only becames 17·6 wt %,
which is 1wt % higher than that of the low-SiO2 melt.
However, this difference in MgO content is still smaller
than expected from the estimated difference in the degree
of melting under isobaric conditions. Slight decreases in
the FeO* contents of the estimated primary melts of the
low- to high-SiO2 groups can also be explained by a combination of effects of pressure and degree of melting.
Therefore, small differences in the MgO contents of the
primary melts as well as FeO* among the three KitaMatsuura basalt groups are consistent with accumulation
of a partial melt generated by polybaric incremental
melting.
Comparison with mantle potential
temperatures estimated in other studies
Kunimi
Low, medium, and high represent low-SiO2, medium-SiO2,
and high-SiO2 groups, respectively.
*K–Ar ages were calculated by regressing separately for
each section to estimate ages for lava flows without age
determination [see details given by Sakuyama et al.
(2009)].
ySamples outside the mixing triangle between C1, C2, and
C3.
The MgO content of a partial melt of peridotite increases with an increase in the degree of melting under isobaric conditions: MgO content generally increases by
42 wt % with an increase of 10% in the degree of melting
(e.g. Hirose & Kushiro, 1993; Baker & Stolper, 1994;
Kushiro, 1996; Pickering-Witter & Johnston, 2000; Schwab
& Johnston, 2001). Therefore, the estimated difference of
410% in the degree of melting between the low- and
high-SiO2 groups may indicate that the MgO contents of
the estimated primary melt compositions should increase
in the order of low-, medium-, and high-SiO2 groups.
However, this expectation is not necessarily the case for
polybaric incremental melting because the MgO content
of the partial melt decreases with decrease of pressure at a
given melting degree. This suggests that the MgO content
of the primary melt does not necessarily increase with an
increase in the degree of melting, especially if melting was
polybaric. Indeed, results of high-pressure experiments on
Putirka (2005) estimated the mantle potential temperature
to be 14508C beneath mid-ocean ridges, which is higher
than other estimates (1280^14008C; McKenzie & Bickle,
1988; Iwamori et al., 1995; Lee et al., 2009). This difference
is mainly due to the fact that the Putirka (2005) study was
based on the most magnesian phenocrysts (Fo# ¼ 91^92)
found in basalts. In contrast, the Fo# of residual olivine
in peridotite assumed in many of the other studies is 89^
90 (e.g. Lee et al., 2009). Lee et al. (2009) obtained a
mantle potential temperature of 1300^14008C beneath
mid-ocean ridges by assuming Fo# ¼ 90. Because the
Fo# of olivine in peridotite increases with the degree of
melting, olivine in peridotite at the lowest pressure and
temperature during adiabatic melting is expected to have
the highest Fo#, if we assume a single episode of peridotite upwelling.
Because the melting degree of the high-SiO2 group is
higher than that of the low-SiO2 group, the Fo# of olivine
in equilibrium with the primary melt of the high-SiO2
group could be higher than 90; however, in this study a
value of 90 was assumed for all three SiO2 groups.
Accordingly, the primary magma for the high-SiO2 group
is most likely to be in equilibrium with olivine having the
highest Fo# in northwestern Kyushu. If we assume an
Fo# of 91 for the residual peridotite of the high-SiO2
group, which is also the highest Fo# of olivine in ultramafic xenoliths from northwestern Kyushu, the mantle potential temperature beneath Kita-Matsuura is estimated
to be 414008C. This estimate is consistent with the result
obtained above for the low-SiO2 group primary melt
assuming Fo# ¼ 90 for the residual olivine in peridotite.
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If one erroneously used Fo# ¼ 91 for the low-SiO2 primary melt, which was generated at the initial melting
stage, the mantle potential temperature would be overestimated at 415008C. Therefore, it is important to constrain
the appropriate Fo# of the olivine in the residual peridotite for each stage of melt segregation for accurate estimation of the mantle potential temperature in an
adiabatically upwelling mantle. Our estimate of the potential temperature beneath Kita-Matsuura is based on the
simple assumption that the low-SiO2 basalt magma was
produced under near-anhydrous conditions and at a lower
degree of melting. Therefore, a simple estimation method
for potential temperature is appropriate based on the olivine fractionation-corrected primary basalt composition, as
noted above. The estimated mantle potential temperature
(14508C) may be slightly higher than in the mid-ocean
ridge setting; however, it is lower than that for ocean
island basalts (Tp416008C; e.g. Putirka, 2005). Our estimate would be reasonable for a back-arc setting and is consistent with the melting regime discussed above.
the lower mantle may have been able to rise through the
rupture, cross the mantle Transition Zone (410^660 km),
and continue upwelling (Fig. 16b). Alternatively, the rupturing of the slab may have triggered upwelling of hotter
mantle from beneath the slab (Fig. 16c) in a similar way
to the slab window model (e.g. Thorkelson & Taylor,
1989). Cenozoic basaltic back-arc volcanism in northwestern Kyushu is aligned ENE^WSW. This distribution may
be caused by the orientation of the rupture beneath
southwestern Japan, although it can also be explained by
near-surface tectonics (Sakuyama et al., 2009). The high
3
He/4He ratio, up to 16 RA, in a mantle xenolith from the
Higashi-Matsuura basalt, which erupted at 3 Ma to the
east of the Kita-Matsuura basalts (Fig. 1), also supports
our model (Sumino et al., 2000). Although we have no age
constraints for the initiation of rupturing or for how long
such a rupture can exist, these models are currently the
most plausible mechanisms to explain the high mantle potential temperatures beneath Kita-Matsuura, which are
higher than those beneath mid-ocean ridges.
Geodynamic implications of a hot mantle
diapir upwelling beneath Kyushu
CONC LUSIONS
Plume models proposed for the eastern margin of the
Eurasian Plate have assumed that the upwelling originates
at either the upper^lower mantle or core^mantle boundary
(e.g. Nakamura et al.,1990).These models, however, may not
entirely rule out the possibility of a shallower origin, as they
were based only on trace element and isotope data.The high
mantle potential temperature estimated for the KitaMatsuura basalt (414508C) suggests that the mantle upwelling responsible for the Kita-Matsuura volcanism originates
from depths that, at the very least, are deeper than the
source mantle of MORB.
Obayashi et al. (2009) observed a discontinuity in the
subducting Pacific Plate beneath Japan; this trends in an
east^west direction from southwestern Japan to the Yellow
Sea at a depth of 300^700 km (Fig. 16a). They interpreted
this discontinuity to be a rupture in the Pacific slab,
which is expected based on the geometry of plate motions.
Obayashi et al. (2006) also observed a low P-wave velocity
region under the Pacific Plate on the oceanward side of
northern Honshu extending from 660 km depth, which
can be observed in the seismic tomography of Huang &
Zhao (2006) (Fig. 16a). They attributed this low-velocity
anomaly to a high-temperature anomaly associated with a
small amount of melt related to hot mantle upwelling
from the lower mantle. If this is the case, the hot upwelling
through the lower mantle may have stopped rising once it
reached the stagnant slab at the 660 km discontinuity,
marking the base of the upper mantle. The upwelling
may then have split and spread horizontally to the east
along the bottom of the slab (Fig. 16b). If the rupture in
the stagnant slab was already present below northwestern
Kyushu at 9 Ma, some of the material upwelling from
As noted above, there are several mechanisms that can account for the temporal and spatial changes in the chemistry, pressure, temperature, degree of melting, and H2O
content of the Kita-Matsuura basalts. Three possible mechanisms can introduce fluid progressively into the hot and
dry upwelling mantle. These include (1) the progressive
introduction of fluid from the deep-seated (mantle
Transition Zone, MTZ) Pacific Plate slab, (2) interaction
of upwelling deep, hot mantle with pre-existing metasomatized, hydrous shallow mantle lithosphere, and (3) entrainment of hydrous wedge mantle material by the upwelling
mantle diapir. The first model is problematic because the
sediment on the stagnant Pacific Plate slab in the MTZ
may have a different composition from that sampled. In
particular, the water content of the slab sediment would
probably be much lower because of dehydration during
subduction beneath the volcanic arc in Japan (Kimura
et al., 2010) and the Marianas (Kelley et al., 2010). Instead,
fluids may have been introduced from subducted serpentinite within the oceanic lithosphere or nominally anhydrous minerals within the MTZ (Richard & Bercovici,
2009). Alternatively, they could be related to melting of Khollandite in subducted slab sediments in the MTZ (Rapp
et al., 2008), which released fluid to metasomatize the
source of the intra-plate basalts above the stagnant Pacific
slab as proposed for northeastern China (Kuritani et al.,
2011). However, such a model would release the water at
depth and does not readily explain the progressive addition of water inferred for the Kita-Matsuura basalts. The
second ‘wet mantle lithosphere’ model maybe a possible
candidate. As the onset of melting and water introduction
begins at a pressure of 3 GPa and ends at 1·5 GPa, as
1118
SAKUYAMA et al.
MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Fig. 16. (a) P-wave velocity anomalies around the Japan Sea at depths from 480 to 550 km (Obayashi et al., 2009) and an east^west vertical
cross-section showing P-wave velocity anomalies along a latitude of 338N (Huang & Zhao, 2006). Red and blue colours denote low and high seismic velocity anomalies, respectively. Green dashed line indicates a plausible rupture of the Pacific Plate. Red dashed line represents 338N. (b)
Schematic illustration of active upwelling from the lower mantle and penetration of the upwelling; (c, d) schematic illustration of passive upwelling from the mantle Transition Zone induced by tearing of the subducted Pacific Plate to generate the Late Miocene basaltic activity in
northwestern Kyushu. Distribution of isotopic end-member components C1, C2, and C3 is shown in (c).
discussed above, the depth range may correlate to the
thickness of the continental lithosphere (Zhu, 2007).
However, the high Tp ¼14508C of the Kita-Matsuura basalts throughout the melting regime precludes involvement
of cool mantle lithosphere (T512008C or less under hydrous conditions). The third model was proposed by
Sakkuyama et al. (2009) to explain the spatio-temporal
variations in chemistry, pressure, and the degree of melting. In this study, we have further clarified the progressive
addition of water and fluid-mobile elements through time.
This new information does not violate the original model
that we proposed in 2009, which still provides the most
plausible mechanism to explain the geochemical data. In
combination with the possibility of deep^hot mantle upwelling, as noted above, we propose again the mechanical
entrainment of metasomatized asthenosphere by a rising
mantle diapir.
AC K N O W L E D G E M E N T S
We are deeply grateful to Yoshiyuki Tatsumi and Hikaru
Iwamori for scientific guidance and constructive
discussions. Sincere thanks are extended to Hiroko
Nagahara for extensive advice. Hideto Yoshida is
also thanked for assistance with the electron microprobe.
We greatly appreciate thoughtful and constructive reviews by Eiichi Takahashi, Jun-Ichi Kimura, Erin Todd,
John Gamble, Marjorie Wilson and two anonymous
reviewers.
FU N DI NG
Part of this work was supported by funds from Ministry of
Education, Culture, Sports, Science and Technology of
Japan (21540495 to M.Y. and 23740398 to T.S.).
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A P P E N D I X A : C O M PA R I S O N O F
T R AC E E L E M E N T
C O N C E N T R AT I O N S A N A LY Z E D
B Y I C P- M S A N D X R F
Results of the comparison are shown in Fig. 17. Rb, Ba,
Th, Nb, and La show good agreement. Sr, Zr, and Y
are higher by 5^10% in the XRF analyses, but still show
reasonable correlation. The discrepancy in Y between
XRF and ICP-MS could have originated in errors in the
standard values (Robinson et al., 1999). Pb shows a weak
correlation between the XRF and ICP-MS data, as the
precision of Pb measured by XRF is much less than
that achieved using ICP-MS, resulting in up to 50%
difference. Elemental ratios used in this study (Zr/Y,
Nb/Y, and Nb/Th) show good correlation and the
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deviations are smaller than 15%. These differences are
much smaller than the variation in each SiO2 group. For
example, if we used ICP-MS data instead of XRF data
for modelling of the Nb/Y^Zr/Y variations, similar to
Sakuyama et al. (2009), the degree of critical melting
based on ICP-MS data would be systematically at most
0·5% lower than that for the XRF data, which is negligibly small.
APPENDIX B:
G E O T H E R M O M E T RY AT H I G H
PRESSU RE
Relationships are observed amongst temperature,
pressure and melt MgO content in high-pressure anhydrous peridotite melting experiments where the melt is
in equilibrium with olivine and pyroxenes (Fig. 18). The
data plotted are taken from the published literature
(Jaques & Green, 1980; Hirose & Kushiro, 1993;
Takahashi et al., 1993; Baker & Stolper, 1994; Hirose &
Kawamura, 1994; Kushiro, 1996; Hirose & Kushiro, 1998;
Robinson et al., 1998; Walter, 1998; Falloon et al., 1999;
Pickering-Witter & Johnston, 2000; Schwab & Johnston,
2001). Regression lines for each pressure calculated using
equation (3) from this study are shown in Fig. 18a and b.
The relationship between melt fraction and temperature
as a function of pressure is shown in Fig. 18c. Regression
lines at each pressure are almost parallel to each other
with good correlation coefficients, most of which
are greater than 0·9, as suggested by Maaloe (2004).
Because the uncertainty of pressure estimation is
0·3 GPa and that of the MgO content of the melt is
1·0 wt %, the uncertainty on the temperature is
308C. Temperatures calculated following the method of
Sugawara (2000) are consistent with our estimates
according to equation (3). The largest and average
differences obtained between equation (3) and Sugawara
(2000) were 118C and 38C, respectively. By using the
MgO content in experimentally produced partial melts
of a peridotite^basalt hybrid source (KG1; Kogiso et al.,
1998) and the melting pressure estimated by normative
projection of the partial melts, we calculated a
temperature from equation (3) to compare with the
actual experimental temperature. The calculated
temperature is lower than the experimental temperature
for KG1 at 3·0 GPa by up to 358C, whereas the
estimated temperatures for pressures lower than 3·0 GPa
are lower than the experimental temperature by only
108C on average and the 1s is 158C. The estimation
method for melting temperature according to equation
(3) is, therefore, applicable to peridotite systems with a
basaltic component less than 50%.
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MANTLE INTERACTION DURING DIAPIRIC UPWELLING
Fig. 17. Comparison of trace element concentrations and ratios analyzed by inductively coupled plasma mass spectrometry (ICP-MS) and
X-ray fluorescence (XRF).
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Fig. 18. (a) Temperature (8C) vs MgO content (wt %) of olivine- and clinopyroxene-saturated melt in anhydrous peridotite melting experiments. Continuous lines are linear regression lines for each pressure. (b) Temperature (8C) vs MgO content (wt %) of experimental melts
plotted in (a) calculated using equation (3). (c) Melt fraction vs temperature at different pressures for the experiments plotted in (a).
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MANTLE INTERACTION DURING DIAPIRIC UPWELLING
A PPEN DI X C: POSSI BI L I T Y FOR
A M E LT I N G M O D E L I N A S Y S T E M
C LOS E D TO I N P U T
Fig. 19. (a) Na2O þ K2O vs SiO2 for the Kita-Matsuura basalts compared with melt compositions from the high-pressure experiments
that were used for testing the geothermometer from this study. (b)
Comparison of temperature predicted by the geothermometer with
that measured in the experiments. (c) Comparison of H2O contents
in the melt predicted by the hygrometer used in this study with those
reported in the experiments.
The presence of HFSE-rich minerals as residual minerals
in the melting system produces a melt with negative
HFSE anomalies. Any titanium-bearing minerals, however, cannot remain as a residual phase in the melting
system at high degrees of melting (7 wt % as estimated
above), as their solubility is high in a basaltic melt
(Ryerson & Watson, 1987). Even if a titanium-bearing
phase remained in the melting system, the magnitude of
the HFSE depletion relative to LILE and LREE rapidly
decreases as the degree of melting increases because
highly incompatible LILE and LREE decrease faster
than HFSE. In the Kita-Matsuura basalts, the extent of
relative depletion in HFSE increases as the abundance
of HFSE decreases. This relationship between HFSE
and the degree of melting does not depend on the
nature of the melting process (e.g. batch or fractional melting). Therefore the possibility of the existence of a
residual HFSE-rich mineral is rejected as being the cause
of the geochemical characteristics of the Kita-Matsuura
basalts.
Melt^rock reactions in the upper mantle may have the
potential to generate HFSE-depleted high-magnesian
andesite and calc-alkaline magma series rocks (e.g.
Kelemen et al., 1993; Grove et al., 2003). To increase HFSE
depletion as their abundance decreases, the assimilated
mass must be greater than the crystallized mass
(Kelemen et al., 1993); however, this process also decreases
the volatile content of the magma as the reaction proceeds. This is not the case for the Kita-Matsuura basalts,
as the estimated water content of the primary magmas
increases as the HFSE content of the magma decreases.
In addition, crystal dissolution from the surrounding
peridotite by a percolating melt, which results in an
increase in the melt mass, would be enhanced along an
inverted geothermal gradient in the mantle, such as in
the mantle wedge above a subducting slab. When the
Kita-Matsuura basaltic magmatism was active, the KitaMatsuura area was situated in the back-arc region, where
the thermal gradient in the upper mantle must be different from that near the wedge corner. Melt^rock reactions,
which may preferentially occur in the wedge corner
where the thermal gradient is inverted, therefore, may
not be effective.
If a series of mafic lithologies with identical compositions to PM1, PM2 and PM3 were present in the peridotite
host (heterogeneous mantle) and if they completely
melted to produce each melt of the Kita-Matsuura basalts,
the SiO2-poor basalts should have formed first followed
by the SiO2-rich basalts within each section in order to
explain the observed temporal major element variations.
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It is, however, difficult for a heterogeneous mantle source
to reproduce such temporal variations, as both the solidus
and liquidus temperatures of a mafic lithology with the
PM3 chemical composition and a higher water content
are estimated to be lower than those of lithologies with a
PM1 chemical composition and lower water content
according to pMELTS calculations. This is inconsistent
with the observed systematic temporal variation of the
least-fractionated basalts of the Kita-Matsuura basalt, as
the source region for each section was homogeneous in
temperature.
A PPEN DI X D: A PPLICA BI LI T Y OF
M AG M A T H E R M OM ET E R A N D
P L A G I O C L A S E ^ M E LT
H Y G RO M E T E R AT L O W
PRESSU RE
We compiled data from experimental studies conducted
under hydrous conditions and applied our method to estimate the temperature and the water content.
Experimental studies that we used are as follows: Grove
et al. (1982, 1990, 1992); Grove & Bryan (1983); Baker &
Eggler (1987); Meen (1987, 1990); Tormey et al. (1987); Grove
& Juster (1989); Juster et al. (1989); Bartels et al. (1991); Fram
& Longhi (1992); Kinzler & Grove (1992); Sisson & Grove
(1993); Snyder et al. (1993); Toplis & Carroll (1995); Yang
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et al. (1996); Muntener et al. (2001); Sano et al. (2001);
Koepke et al. (2004); Sano & Yamashita (2004); Berndt
et al. (2005); Takagi et al. (2005); Feig et al. (2006, 2010);
Pertermann & Lundstrom (2006); Scoates et al. (2006);
Hamada & Fujii (2007, 2008); Freise et al. (2009); Tatsumi
& Suzuki (2009); Parman et al. (2011); Honma (2012). The
experimental pressure and temperature and the water contents in the melt varied from 0 to 1·5 GPa, from 943 to
12958C, and from 0 to 6·0 wt %, respectively. The range of
the experimental melt compositions was subalkalic to alkalic (Fig. 19a) with SiO2446·9 wt %. However, the minimum value of the SiO2 content of the experimental melts
under hydrous conditions was 48·5 wt %, which is 1wt
% higher than that of the Kita-Matsuura basalts (47·5 wt
%). To check how reliably this method can be applied to
the Kita-Matsuura basalts, the experiments were filtered
by An# [¼ 100 Ca/(Ca þ Na)] from 60 to 90 and SiO2
contents less than 60 wt %.
Calculated temperature and H2O contents in the melt
were compared with experimental temperature and H2O
content in melt (Fig. 19b and c). Temperature estimates
show a good agreement with experimental temperatures
(Fig. 19b); the regression line is y ¼ 0·9951x and
R2 ¼ 0·8975. In contrast, the estimated H2O contents in
the melt showed a slight shift to values lower than the
experimental H2O contents (y ¼ 0·873x), although a fairly
good correlation was obtained (R2 ¼ 0·873).
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