Deep Crystallization Differentiation of Arc Tholeiite Basalt Magmas

Deep Crystallization Differentiation of ArcTholeiite Basalt Magmas from NorthernHonshu Arc, Japan

TSUKASA OHBA1*, KAZUHIDE MATSUOKA2,YASUYUKI KIMURA3,HIROMASA ISHIKAWA4 AND HIROKAZU FUJIMAKI5

1FACULTY OF ENGINEERING AND RESOURCE SCIENCE, AKITA UNIVERSITY, 1-1 TEGATAGAKUEN-MACHI,

AKITA, JAPAN2MITSUBISHI MATERIAL TECHNO CORPORATION, 3-7-13, TOYO, KOTO WARD, TOKYO, JAPAN3NIPPON TELEGRAPH AND TELEPHONE EAST CORPORATION, 3-19-2, NISHISHINJUKU, SHINJUKU KU, TOKYO, JAPAN4SUMIKO CONSULTANTS CO., LTD, 2-9-7, IKENOHATA, TAITO KU, TOKYO, JAPAN5GRADUATE SCHOOL OF SCIENCE, TOHOKU UNIVERSITY, 6-3, ARAMAKI AZA AOBA, AOBA KU, SENDAI, JAPAN

RECEIVED MAY 25, 2007; ACCEPTED MAY 1, 2009

The depth of crystallization differentiation was investigated for arc

tholeiite basalts from the Takada Ohdake and Nishimori cones,

within the volcanic front of the northern Honshu volcanic arc,

Japan, based on the petrography of crystal-bearing melt inclusions,

whole-rock trace element chemistry, and the simulation of fractional

crystallization with MELTS. Olivine-hosted melt inclusions contain

aluminous clinopyroxene, spinel, plagioclase (An60^65), and rare

garnet, suggesting melt-entrapment under garnet-granulite facies

conditions. Removal of the observed phenocrysts (olivine, orthopyrox-

ene, and anorthitic plagioclase) cannot account for either the major

or trace element variations within the basalts. However, the major

element variations are typical of a tholeiitic differentiation trend,

showing an increase in Fe/Mg ratio with increasing SiO2. The

results of trace element modelling, assuming Rayleigh fractionation,

indicate fractionation of clinopyroxene, plagioclase, and spinel.The

estimated fractionated mineral assemblage is consistent with that

in the melt inclusions except for the presence of garnet.The occurrence

of garnet in some of the more differentiated andesitic melt inclu-

sions implies its crystallization within the inclusions during the

later stages of cooling. Fractional crystallization modelling using

MELTS closely approximates the major element variations at

10 kbar under anhydrous conditions. The MELTS modelling is

consistent both with the melt-inclusion mineralogy and with the

trace element modelling. An estimated depth of crystallization

of 34 km (�10 kbar) corresponds to the seismic Moho in the

region (35^36 km). Our data also suggest that the relatively anhy-

drous magmas that were emplaced at Moho depths became hydrous

concurrent with differentiation, implying incorporation of water

from the surrounding crustal rocks.

KEY WORDS: aluminous clinopyroxene; arc basalt; crystal-bearing melt

inclusions; crystallization-differentiation; garnet; northern Honshu

I NTRODUCTIONNorthern Honshu is a well known for its subduction-related volcanic activity for which many petrological andgeochemical data have been accumulated over the past25 years (e.g. Ishikawa et al., 1984; Nakagawa et al., 1985;Fujinawa, 1988, 1990; Hunter & Blake, 1995; Ohba &Umeda, 1999; Ban & Yamamoto, 2002; Toya et al., 2005;Kimura & Yoshida, 2006; Hirotani & Ban, 2006; Ohbaet al., 2007). Two magma series, a calc-alkaline and a tho-leiite series, have been identified in many of the volcanoesalong the volcanic front. As a rule of thumb, chemicalvariations in the tholeiite series of magmas can beaccounted for by crystallization-differentiation processes.Conflicting ideas about the nature of the fractionatingminerals have been proposed for individual volcanoes.

*Corresponding author. E-mail: [email protected]

� The Author 2009. Published by Oxford University Press. Allrights reserved. For Permissions, please e-mail: [email protected]

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Fractionation of phenocrysts was demonstrated for thetholeiitic rock series from Adatara volcano based on exten-sive petrological data (Fujinawa, 1988, 1990). In contrast,some researchers have proposed that simple fractionationof the observed phenocryst minerals is insufficient toaccount for the chemical variation of the arc tholeiitesuite from Iwate, Akita Komagatake, and Nasu (Ishikawaet al., 1984; Nakagawa, 1985; Ban, 1991). Nevertheless, anumber of workers have invoked single-phase fractionationof olivine, known as the olivine maximum fractionationmodel, in studies of regional arc volcanism (Tatsumi et al.,1983; Yamashita & Tatsumi, 1994; Yamashita et al., 1996;Tamura et al., 2000), although this model has not been ade-quately evaluated.In this study our interest is in the depth at which the

parental magmas were emplaced and differentiated.We believe that determination of the depth of crystal-lization will resolve issues concerning the nature of thefractionating mineral assemblage. Evolutionary trends inmagma composition by crystallization differentiation aregoverned by the phase relationships between the crystalliz-ing minerals, and accordingly, by the physical conditionsthat determine the phase equilibria. Crystallization depthand pressure play a controversial role in constrainingphase relations. Recently, crystallization differentiation inthe deep crust has become a focus for discussion. In consid-eration of the genesis of intermediate and silicic magmasat convergent margins, emplacement of basalt magmasinto the deep crust has been assumed in a number of theo-retical models (e.g. Annen & Sparks, 2002; Annen et al.,2006), which involve repeated basalt intrusion into thedeep crust. Davidson et al. (2007) validated these modelsthrough a review of rare earth element data from variousarc magmas. The salient conclusion of that study was thatfractionation of cryptic amphibole at depth in the arccrust plays an important role in the petrogenesis of evolvedarc magmas.The concept of cryptic mineral fractionation vastly

improves our understanding of the differentiation ofmagmas. Large phenocrysts are often regarded as mineralsentrained from a crustal magma chamber in which crystal-lization-differentiation has occurred. However, no basisexists for such an assumption because large crystals can beformed by rapid crystal growth promoted, for example,by decompression-induced degassing as magmas risetowards the surface (Brophy et al., 1999; Kuritani, 1999;Blundy & Cashman, 2001, 2005; Annen et al., 2006) or byrapid cooling during magma mingling (Ohba et al., 2007).High-pressure mineral phases might disappear duringmagma ascent if they are unstable at lower pressures.Therefore, visible phenocrysts are not necessarily identicalto the fractionated minerals in evolving crustal magmachambers. Because of this classic petrographic percep-tion, the equilibrium phase relations determined from

experiments might not have been applied properly instudies of erupted magmas, leading to incorrect estimatesof crystallization conditions.In this study we examine two subjects: (1) the origin of

phenocrysts in arc magmas; (2) the physical conditions ofcrystallization differentiation based on petrographic andbulk-rock geochemical and mineral chemical data fortwo basaltic cones in northern Honshu, Japan (TakadaOhdake and Nishimori). Our goal is to determine magmaemplacement depths in the crust. Our approach involvespetrography, trace element modelling, and major elementmodelling by MELTS simulation. The petrographicapproach specifically examines the mineralogy ofphenocryst-hosted melt inclusions that might retain evi-dence of higher pressure magma differentiation conditions.Our target fields are two basaltic cones: Takada Ohdakeand Nishimori. Rocks from these basaltic cones are suit-able for examining crystallization differentiation of primi-tive parental magmas beneath the arc for a number ofreasons. The cones mostly comprise tholeiitic rocks, whoserange of chemical variation is considered to result fromcrystallization differentiation. In the Japan arc, magmasystems beneath most of the volcanoes are complex; tholeii-tic rocks are intimately associated with abundant calc-alkaline magmatic rocks that are usually interpreted asoriginating from processes of magma mixing, crustalassimilation, or lower crustal melting. Two large cones inthe region, Iwate and Akita Komagatake, are composedexclusively of tholeiitic rocks, which may have experiencedmultistage magma evolution during the construction ofthese large and complex volcanic edifices. In contrast,Takada Ohdake and Nishimori are small, short-lived vol-canic cones.We therefore consider that multistage magmaevolution is less likely beneath these and that they presentmuch simpler cases for interpretation of crystallizationdifferentiation.

GEOLOGICAL BACKGROUNDThe northern Honshu volcanic arc is located in a conver-gent margin where the Pacific Plate is subducting beneaththe northern part of Honshu. Along the arc volcanicfront, lavas of medium-K calc-alkaline andesite and low-K tholeiitic basalt are dominant along with caldera-forming rhyolite ignimbrites. These andesite and basaltlavas constitute stratovolcanoes that are mostly compoundcomposite cones. The stratovolcanoes and calderas areassembled to form large volcanic clusters. Seven volcanicclusters have been identified in northern Honshu: Osore^Hiuchi, Towada^Hakkoda, Sengan, Kurikoma, Zao,Bandai, and Nasu (Fig. 1).Takada Ohdake (408390110 0N, 14085402900E) is a small

symmetric stratocone located in the northeast part of theTowada^Hakkoda cluster. The cone is about 3 km in diam-eter and some 600m high and consists of low-K tholeiite

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basalt lavas and associated scoria fall deposits. TakadaOhdake has been regarded as forming on part of NorthHakkoda, a compound stratovolcano comprising 10 smallstratocones, associated lavas, and debris avalanche depos-its. Other cones in North Hakkoda are composed mainlyof andesite and dacite. Takada Ohdake is an extinctvolcano that was active from 0�3 to 0�4Ma (Kudo et al.,2004). Rocks were sampled along a southern trail, northerngullies, and a gorge at the eastern base. Piles of thinbasalt lavas are exposed successively along the easterngorge which developed on a topographic low throughwhich lava flowed.

Nishimori (39858’39’’N, 1408N56’25’’E) is a small strato-cone consisting of low-K tholeiite basaltic lavas and scoriafall deposits. The cone is located at the northeastern end ofthe Hachimantai compound stratovolcano, a member ofthe Sengan volcano cluster. The cone shape is obscured byuneven basement relief, by partial dissection and bya covering of younger deposits.The cone overlies the depos-its of older andesite^dacite cones (Ebisumori, Chausudake,Appidake) on thewest; its eastern flank is dissected and cov-ered by the younger basaltic composite cone of Maemori.Rocks from the cone have not been dated, but the activityage can be determined from stratigraphic relations.

Fig. 1. Maps of (a) the Japanese Islands and (b) the northern Honshu volcanic arc showing the locations of volcanic clusters (shaded areas).Takada Ohdake and Nishimori are located respectively in the Towada^Hakkoda and Sengan volcanic clusters. The topography of the TakadaOhdake and Nishimori cones are shown in (c) and (d), respectively. Heavily shaded areas are the eruptive products of Takada Ohdake andNishimori. �, sampling locations.

OHBA et al. DEEP CRYSTALLIZATION DIFFERENTIATION

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Its underlying andesite is 0�4^0�7Myr old (KÂr and TL;Ohba & Umeda, 1999; Takashima et al., 2001). The age ofthe overlying basalt is 0�1^0�2Ma (Doi, 2000). Therefore,the active period of Nishimori is inferred as 0�4^0�2Ma.Rocks were sampled from outcrops near the summit andalong a gully at the northern base.

ANALYT ICAL METHODSA thin section and rock powder were prepared from eachsample. From selected samples, polished thin sectionswere prepared for electron microprobe analysis. Sampleswere powdered by grinding in an agate mill for whole-rock analysis by X-ray fluorescence spectrometry (XRF)and inductively coupled plasma mass spectrometry(ICP-MS). Back-scattered electron images (BSEI) werecaptured using a scanning electron microscope (SEM:JSM-5410; JEOL) atTohoku University. Microprobe analy-ses of minerals were performed by energy dispersive spec-troscopy (EDS: ISIS Link; Oxford Co. Ltd.) attachedto an SEM and a wavelength-dispersive spectroscopy(WDS) electron probe microanalyzer (EPMA: JWM-8800; JEOL). Analyses of major elements and some traceelements (Sr, Rb, V, Ni) were carried out by XRF(RIX2100; Rigaku Corp.). The glass-bead method(sample:LiB4O7¼1:5) was used. Calibration curves weredetermined using GSJ Igneous Rock Series standards andcorrected by empirical matrix factors. Because Rb abun-dances are low in the samples, but necessary for meaning-ful discussion, the calibration curve was determined usingstandard samples with low Rb contents (520 ppm).The accuracy of the calibration curve for Rb is c. 0�3 ppm;relative errors are less than 5%. Accuracies are less than6 ppm for Ni, Sr, and c. 10 ppm for Cr and V. Rare earthelements (REE) and some trace elements (Zr, Hf, Ta, andTh) were analyzed at Tohoku University by ICP-MS(Elan6000; Perkin Elmer Co.). Calibration with dilutedHNO3 solutions was used. The 1:5 glass bead used in XRFanalysis was crushed to a powder and dissolved in aHNO3 solution to a dilution factor of 1/1000. The GSJsamples (JG-2 and JA-3) were used as standards; thecalibration curves determined using these samples wereverified using other GSJ Igneous Rock Series standards(see Supplementary Data 1, available for downloadingat http://www.petrology.oxfordjournals.org). Whole-rockcompositions are listed inTable 1, and the compositions ofphenocryst and minerals in melt inclusions are listedrespectively inTables 2 and 3.

PETROGRAPHY OF THESAMPLESPhenocryst mineralogyAll samples are porphyritic and contain abundant pheno-crysts of plagioclase and mafic minerals. Four types of

mafic phenocryst assemblage occur in Takada Ohdake:(1) olivine; (2) olivineþ orthopyroxene; (3) olivineþorthopyroxeneþ augite; (4) orthopyroxeneþ augite. Twoof these assemblages occur at Nishimori (olivineþ ortho-pyroxeneþ augite and orthopyroxeneþ augite). AtTakada Ohdake, the mafic phenocryst assemblages areroughly correlated with bulk-rock MgO contents (Table 1).Mineral assemblages in the most magnesian samples areol and olþopx, those of intermediate composition containolþopxþcpx; the least magnesian samples containopxþcpx.Plagioclase phenocrysts have highly calcic (An88^95)

cores and less calcic rims (An70) in most samples fromboth cones, irrespective of the degree of differentiation(i.e. whole-rock Na2O and MgO content). Exceptions areplagioclase phenocrysts in andesite samples (453% SiO2),of which the plagioclase core compositions range broadlyfrom An65 to An90. Mafic mineral compositions varybroadly in terms of MgO/FeO corresponding to theMgO/FeO variation of the whole-rocks. The olivine com-position varies from Fo81 in a magnesian sample and Fo47in a differentiated sample. Even the most magnesianminerals (olivine and orthopyroxene in Nishi-s3 and5072004) have much lower Mg/(MgþFe) than mantle-equilibrated compositions, implying that even the mostmagnesian basalts have experienced some crystallizationdifferentiation. Phenocryst clinopyroxene is augite incomposition with low Al2O3 (54%) and Na2O (50�15%)contents. Orthopyroxene is typically hypersthene withlowAl2O3 (52%).

Minerals in melt inclusionsCrystal-bearing melt inclusions occur in large olivinephenocrysts (Fig. 2) in a pyroxene-free olivine basaltfrom Takada Ohdake (sample number 5072004) and ina clinopyroxeneôrthopyroxeneôlivine basalt fromNishimori (Nishi-s3, samples HM0703A, and HM0701D).The inclusions are 4^200 mm, partly occupied by vapourbubbles, and contain crystals of various sizes of abundantaluminous clinopyroxene, common spinel, plagioclase,orthopyroxene, and rare garnet. Garnet occurs only inthe olivine basalt from Takada Ohdake (5072004); it hasnot been identified in melt inclusions in any of the othersamples. Orthopyroxene is found in only one sample(HM0703A) from Nishimori. Aluminous clinopyroxene isthe commonest mineral observed in the melt inclusions.Alumina contents vary widely, even in a single sample,from 8 to 17wt % (Table 3). The clinopyroxene is poor inNa2O (50�4%), and is, therefore, poor in the omphacitecomponent. The aluminium ions are distributed in bothfour- and six-fold coordinated sites, and the pyroxene isrich in the M2þAlVIAlIVSiO6 (M2þ

¼Ca2þ, Fe2þ, Mg2þ)rather than the M2þFe3þAlSiO6 component. Similarclinopyroxene has been reported only rarely for highP^T metamorphic rocks (e.g. Ro« tzler & Romer, 2001).


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Table 1: Whole-rock analyses of rocks fromTakada Ohdake and Nishimori

Cone: Takada Ohdake

Sample no.: 5072003 5072004 HK0701 5072007 5072008 5072105N HK0702 5072103S 5072102S 5072104S 5072105S 5072107S 5072101N 5072103N 5072009 5072102N

Latitude (N): 40 39’56" 40 39’56" 40 39’49" 40 39’49" 40 39’45" 40 39’56" 40 39’35" 40 38’48" 40 38’41" 40 38’52" 40 38’58" 40 39’05" 40 39’56" 40 39’45" 40 39’39" 40 39’46"

Longitude (E): 140 57’00" 140 56’24" 140 56’48" 140 56’43" 140 56’37" 140 54’53" 140 56’54" 140 54’49" 140 54’56" 140 54’43" 140 54’38" 140 54’27" 140 55’03" 140 54’36" 140 56’35" 140 54’35"

Mafic minerals: (opx) ol ol (opx) ol (opx) ol (opx) ol (cpx (opx) ol (ol) ol (ol) (ol) cpx opx (ol) (ol) ol (ol)

opx) ol cpx opx cpx opx cpx opx cpx opx cpx opx cpx opx cpx opx cpx opx

wt %

SiO2 48�60 49�45 49�77 49�68 49�84 50�70 50�01 52�46 53�07 52�88 53�81 53�60 51�69 52�96 52�00 53�08

TiO2 0�66 0�75 0�81 0�75 0�81 0�80 0�83 0�73 0�74 0�75 0�76 0�75 0�82 0�87 0�84 0�89

Al2O3 16�42 16�99 17�19 17�91 18�36 17�95 17�91 18�00 17�90 17�91 17�60 17�87 18�95 17�90 19�21 17�82

Fe2O3t 11�81 11�83 11�59 11�56 11�52 11�02 11�30 10�31 10�25 10�10 10�03 9�90 10�51 11�01 10�46 10�92

MnO 0�181 0�183 0�184 0�181 0�181 0�177 0�182 0�170 0�168 0�166 0�165 0�164 0�169 0�182 0�170 0�183

MgO 10�04 8�77 7�69 7�31 6�72 6�55 6�54 5�82 5�81 5�61 5�52 5�50 5�33 4�98 4�91 4�88

CaO 9�24 9�15 9�91 10�15 9�89 10�20 10�35 9�62 9�55 9�70 9�11 9�42 10�05 9�50 9�91 9�47

Na2O 1�86 1�84 2�10 2�03 2�06 2�20 2�14 2�23 2�29 2�36 2�34 2�27 2�52 2�57 2�42 2�63

K2O 0�14 0�14 0�16 0�15 0�15 0�21 0�15 0�38 0�38 0�39 0�45 0�40 0�19 0�24 0�16 0�26

P2O5 0�078 0�073 0�087 0�087 0�087 0�100 0�091 0�090 0�078 0�095 0�075 0�080 0�101 0�100 0�088 0�104

H2O (–) 0�44 0�58 0�28 0�35 0�72 0�52 0�59 0�51 0�17 0�40 0�78 0�62 0�66 0�29 0�39 0�56

LOI –0�40 0�15 0�44 –0�60 0�41 –0�27 0�09 –0�49 –0�13 –0�47 0�07 –0�23 –0�33 –0�42 –0�08 –0�39

FeOt/MgO 1�06 1�21 1�36 1�42 1�54 1�51 1�56 1�60 1�59 1�62 1�63 1�62 1�77 1�99 1�92 2�01

Trace elements (XRF) (ppm)

V 241 253 260 268 277 268 265 238 248 240 238 235 243 269 233 287

Cr 197 187 163 119 121 80 112 69 66 63 62 61 92 50 51 52

Ni 127 92 71 59 52 39 44 36 31 32 29 29 38 19 25 18

Rb 5�4 4�5 4�6 5�4 4�0 6�9 3�2 10�1 10�5 10�5 11�9 10�5 5�1 6�0 4�0 6�3

Sr 282 280 303 299 301 296 313 287 290 297 284 295 335 291 351 289

Ba 83 63 67 69 84 91 69 121 130 129 144 140 83 104 112 106

Trace elements (ICP-MS) (ppm)

Y 13 14 16 15 15 17 16 18 20 15 19 20 18 21 18 21

Zr 28 31 32 31 33 40 33 50 53 42 56 55 38 45 44 47

Nb 0�94 1�05 1�39 1�05 1�07 1�28 1�50 1�36 1�45 1�13 1�48 1�53 1�28 1�34 1�44 1�36

Cs 0�06 0�00 0�32 0�02 0�00 0�39 0�22 0�52 0�62 0�38 0�78 0�57 0�26 0�36 0�03 0�44

La 2�45 2�64 2�76 2�61 2�75 3�36 2�86 3�92 4�15 3�29 4�40 4�25 3�40 3�47 3�81 3�49

Ce 6�50 6�73 7�28 7�03 7�19 8�86 7�16 9�81 10�73 8�42 11�21 11�13 8�56 9�19 10�30 9�36

Pr 0�97 1�03 1�11 1�05 1�09 1�32 1�17 1�43 1�56 1�22 1�64 1�62 1�32 1�37 1�48 1�41

Nd 4�98 5�36 5�82 5�55 5�82 6�67 5�94 7�08 7�53 6�01 7�80 7�85 6�70 7�14 7�45 7�35

Sm 1�56 1�67 1�86 1�73 1�77 2�04 1�88 2�11 2�28 1�80 2�33 2�34 2�12 2�31 2�26 2�35

Eu 0�64 0�68 0�75 0�69 0�72 0�74 0�77 0�73 0�76 0�59 0�72 0�76 0�80 0�86 0�84 0�87

Gd 1�69 1�84 2�03 1�84 1�94 2�25 2�07 2�39 2�54 1�95 2�53 2�57 2�37 2�72 2�40 2�72

Tb 0�36 0�38 0�42 0�40 0�41 0�46 0�44 0�49 0�52 0�41 0�51 0�53 0�48 0�55 0�51 0�54

Dy 1�96 2�11 2�35 2�17 2�25 2�57 2�40 2�71 2�84 2�24 2�89 2�94 2�63 3�10 2�78 3�10

Ho 0�42 0�46 0�51 0�47 0�49 0�55 0�51 0�59 0�60 0�48 0�62 0�63 0�57 0�66 0�58 0�67

Er 1�28 1�40 1�53 1�45 1�48 1�71 1�64 1�84 1�94 1�49 1�95 1�95 1�73 2�03 1�79 2�11

Tm 0�20 0�21 0�22 0�22 0�22 0�26 0�23 0�27 0�29 0�23 0�30 0�30 0�26 0�32 0�27 0�31

Yb 1�37 1�50 1�62 1�50 1�59 1�79 1�59 1�97 2�06 1�63 2�11 2�15 1�85 2�26 1�89 2�27

Lu 0�21 0�23 0�24 0�24 0�24 0�28 0�25 0�30 0�31 0�25 0�32 0�33 0�29 0�35 0�29 0�34

Hf 0�82 0�94 1�13 0�91 0�93 1�16 1�08 1�48 1�51 1�23 1�67 1�68 1�06 1�35 1�26 1�31

Ta 0�07 0�08 0�15 0�07 0�07 0�09 0�14 0�10 0�11 0�09 0�10 0�12 0�08 0�14 0�09 0�09

Th 0�32 0�35 0�39 0�36 0�37 0�55 0�38 1�05 1�06 0�90 1�30 1�29 0�34 0�52 0�49 0�53

Sc 32 35 32 34 37 35 33 33 34 26 31 31 33 39 34 39

(continued)


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Table 1: Continued

Cone: Takada

Ohdake

Nishimori

Sample no.: HK0706 HK0704 nishi-s3 Nishi-t1 nishi-s4 nishi-t2 nishi-s1 HM0701A nishi-s2 HM0701E HM0703A HM0701B HM0702 asn3 asn1

Latitude (N): 40 39’19" 40 39’27 39 58’43" 39 58’36" 39 58’43" 39 58’36" 39 58’43" 39 58’43" 39 58’43" 39 58’43" 39 58’43" 39 58’43" 39 58’38" 39 59’00" 39 59’15"

Longitude (E): 140 56’28" 140 56’34" 140 56’36" 140 56’41" 140 56’36" 140 56’41" 140 56’48" 140 56’48" 140 56’48" 140 56’48" 140 56’36" 140 57’36" 140 56’46" 140 55’51" 140 55’47"

Mafic minerals: cpx opx cpx ol opx (cpx) ol opx (cpx) opx ol (cpx) opx ol (cpx) opx ol (cpx) opx ol cpx opx ol (cpx opx) ol opx cpx ol (opx cpx) ol opx cpx ol (cpx opx) ol (cpx opx) ol ol opx cpx

wt %

SiO2 51�46 52�43 51�34 51�42 51�12 51�78 51�11 50�94 50�84 51�66 50�16 50�72 50�91 50�01 52�16

TiO2 0�84 0�84 0�58 0�62 0�61 0�62 0�63 0�60 0�63 0�59 0�59 0�58 0�69 0�69 0�76

Al2O3 19�20 18�76 15�38 15�97 15�85 16�08 16�21 16�10 16�27 16�10 16�16 16�29 17�67 18�21 17�57

Fe2O3t 10�39 10�22 11�06 10�72 11�06 10�57 10�84 10�31 11�22 9�87 10�48 10�28 10�10 10�98 10�39

MnO 0�172 0�170 0�191 0�195 0�194 0�189 0�192 0�185 0�193 0�182 0�188 0�182 0�174 0�187 0�175

MgO 4�71 4�56 7�93 7�47 7�45 7�45 7�44 7�09 7�04 6�93 6�91 6�89 6�42 5�72 5�55

CaO 10�04 9�78 10�32 10�28 10�54 10�43 10�25 10�42 10�61 10�13 10�31 10�40 10�64 11�29 10�09

Na2O 2�43 2�44 1�65 1�63 1�72 1�74 1�66 1�76 1�70 1�71 1�72 1�72 1�92 1�77 2�18

K2O 0�13 0�27 0�29 0�27 0�29 0�25 0�26 0�26 0�26 0�28 0�26 0�24 0�18 0�17 0�25

P2O5 0�117 0�116 0�068 0�069 0�070 0�070 0�071 0�070 0�071 0�068 0�075 0�068 0�076 0�072 0�087

H2O (–) 0�63 0�42 0�79 1�43 0�95 0�98 1�22 0�21 0�88 4�74 0�21 0�63 0�18 0�50 0�42

LOI 0�20 0�21 �0�30 �0�15 0�36 –0�18 –0�29 0�19 –0�68 0�51 0�19 0�06 0�34 –0�76 –0�45

FeOt/MgO 1�99 2�02 1�25 1�29 1�34 1�28 1�31 1�31 1�44 1�28 1�37 1�34 1�42 1�73 1�69


V 244 241 248 249 254 257 250 238 262 231 237 234 260 272 276

Cr 48 44 256 240 250 207 212 203 198 208 185 183 122 110 78

Ni 23 23 69 64 62 62 61 53 57 57 52 51 38 27 25

Rb 2�5 6�3 10�1 7�0 10�4 6�1 8�8 9�0 7�7 5�7 6�9 6�9 2�0 6�8 6�3

Sr 353 348 219 227 223 229 234 225 230 225 224 224 281 281 291

Ba 100 130 82 91 87 90 85 82 81 92 80 79 74 64 105


Y 19 21 14 16 14 15 14 14 14 14 13 13 13 13 19

Zr 44 47 35 55 35 40 34 32 36 46 32 31 29 28 71

Nb 1�90 2�07 1�91 2�18 1�89 1�99 1�78 2�25 1�80 2�32 2�12 2�12 1�82 1�33 2�58

Cs 0�18 0�37 0�08 0�13 0�16 0�07 0�16 0�15 0�08 0�13 0�10 0�11 0�06 0�06 0�16

La 3�79 4�19 4�10 3�76 3�96 4�12 3�68 3�85 3�66 4�60 3�43 3�38 2�90 2�61 4�80

Ce 9�95 10�57 9�30 9�18 9�50 9�62 8�97 8�82 8�71 9�60 8�33 7�95 6�92 6�65 10�91

Pr 1�50 1�62 1�18 1�21 1�22 1�28 1�19 1�20 1�13 1�37 1�11 1�08 1�05 0�93 1�49

Nd 7�63 8�19 5�67 5�70 5�69 5�94 5�54 5�75 5�48 5�17 5�23 5�20 5�96 4�65 6�96

Sm 2�32 2�50 1�58 1�68 1�63 1�66 1�56 1�66 1�55 1�59 1�52 1�46 1�68 1�42 1�97

Eu 0�91 0�93 0�63 0�65 0�64 0�66 0�63 0�64 0�64 0�67 0�59 0�59 0�68 0�59 0�74

Gd 2�44 2�60 1�70 1�80 1�78 1�84 1�70 1�77 1�71 1�69 1�60 1�58 1�79 1�56 2�13

Tb 0�50 0�55 0�34 0�38 0�35 0�36 0�35 0�37 0�35 0�36 0�34 0�33 0�36 0�33 0�43

Dy 2�80 2�96 1�97 2�13 1�97 2�06 1�98 2�08 2�00 1�87 1�87 1�89 2�02 1�79 2�50

Ho 0�59 0�64 0�42 0�48 0�43 0�44 0�43 0�43 0�43 0�42 0�41 0�40 0�44 0�31 0�56

Er 1�75 1�92 1�30 1�50 1�33 1�35 1�32 1�38 1�34 1�31 1�26 1�23 1�31 1�17 1�78

Tm 0�26 0�28 0�20 0�23 0�20 0�20 0�19 0�20 0�20 0�18 0�18 0�18 0�19 0�18 0�28

Yb 1�82 1�95 1�37 1�63 1�40 1�47 1�42 1�43 1�45 1�36 1�34 1�33 1�39 1�25 1�99

Lu 0�29 0�31 0�21 0�26 0�21 0�22 0�21 0�22 0�22 0�21 0�21 0�21 0�21 0�18 0�30

Hf 1�21 1�41 0�89 1�14 0�90 0�97 0�89 0�88 0�96 0�81 0�82 0�81 0�93 0�75 1�58

Ta 0�17 0�18 0�16 0�17 0�12 0�15 0�12 0�20 0�12 0�17 0�16 0�18 0�21 0�08 0�22

Th 0�47 0�53 0�90 1�04 0�84 0�88 0�78 0�79 0�80 0�47 0�78 0�74 1�03 0�38 1�17

Sc 30 31 36 37 37 37 37 33 36 33 32 33 32 37 37

(continued)


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Table 1: Continued

Cone: Nishimori

Sample no.: asn6 iwht2 iwht3 iwht1 toak1 asn2

Latitude (N): 39 58’53" 39 59’26" 39 59’26" 39 59’26" 39 59’23" 39 59’09"

Longitude (E): 140 55’55" 140 55’31" 140 55’26" 140 55’35" 140 54’52" 140 55’49"

Mafic minerals: cpx opx cpx opx (ol) cpx op cpx opx ol cpx opx

wt %

SiO2 52�52 52�91 53�60 55�17 50�54 57�24

TiO2 0�82 0�90 0�89 0�84 0�80 0�70

Al2O3 18�22 17�34 18�36 16�58 19�49 16�81

Fe2O3t 10�59 11�92 10�55 10�78 10�62 8�51

MnO 0�181 0�200 0�187 0�174 0�175 0�140

MgO 5�16 4�85 4�43 4�27 4�15 4�02

CaO 9�01 8�30 7�49 8�25 10�92 7�58

Na2O 2�14 2�08 1�93 2�35 2�21 2�32

K2O 0�18 0�28 0�30 0�66 0�16 1�02

P2O5 0�097 0�112 0�120 0�108 0�084 0�098

H2O (-) 1�57 1�54 2�46 0�77 1�20 1�21

LOI 0�47 �0�04 1�62 -0�66 �0�42 0�65

FeOt/MgO 1�85 2�21 2�15 2�27 2�30 1�90


V 271 332 295 283 278 203

Cr 68 29 44 49 45 65

Ni 23 15 16 18 18 25

Rb 4�9 6�6 8�9 18�3 3�6 29�8

Sr 282 276 282 261 287 243

Ba 130 122 234 189 93 295


Y 20 18 25 22 19 14

Zr 54 47 93 83 49 36

Nb 2�36 2�89 3�65 3�35 1�84 1�85

Cs 0�08 0�10 0�36 0�75 0�04 0�07

La 5�86 5�01 7�94 7�00 3�86 3�99

Ce 14�01 11�86 18�76 16�39 9�22 9�34

Pr 1�84 1�59 2�46 2�20 1�35 1�23

Nd 8�53 7�51 11�18 10�05 6�75 5�75

Sm 2�32 2�13 2�96 2�68 1�98 1�62

Eu 0�88 0�85 0�96 0�89 0�81 0�64

Gd 2�45 2�29 3�06 2�78 2�25 1�77

Tb 0�49 0�46 0�62 0�55 0�46 0�36

Dy 2�75 2�49 3�45 3�11 2�67 2�03

Ho 0�58 0�56 0�73 0�65 0�57 0�43

Er 1�80 1�68 2�25 2�04 1�80 1�33

Tm 0�26 0�25 0�33 0�31 0�27 0�21

Yb 1�87 1�75 2�33 2�13 1�90 1�45

Lu 0�28 0�27 0�35 0�31 0�29 0�21

Hf 1�37 1�20 2�32 2�08 1�28 0�92

Ta 0�00 0�18 0�23 0�21 0�12 0�12

Th 1�21 0�94 2�41 2�13 0�60 0�85

Sc 36 39 37 34 34 38

Abbreviations of mafic minerals: ol, olivine; cpx, clinopyroxene (augite); opx, orthopyroxene. Orders of minerals indicateascending order in modal volume. Minerals in parentheses are present in trace amount.


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Orthopyroxene in the melt inclusions is also alumina-rich,containing c. 6 wt % Al2O3. This alumina content isunusually high for volcanic rocks, and still higher thanthat of orthopyroxene in mantle-derived rocks. These pyr-oxenes differ compositionally from phenocryst pyroxenesin the same samples in terms of their alumina contents(c. 1^4wt % in augite phenocrysts and 1^2% in ortho-pyroxene phenocrysts). Garnet occurs rarely as very smallgrains (55 mm) contained in melt inclusions in the olivinebasalt from Takada Ohdake (5072004), but has not beenidentified in any of the other samples. The garnet crystalsare contained within spinel and aluminous clinopyroxenecrystals in the olivine-hosted melt inclusions. Becausemost grains are very small, accurate microprobe analysiswas impossible.We performed only two analyses, althoughwe identified several garnet grains by semi-quantitativemicroprobe analysis. The garnet is rich in almandine andgrossular components. The pyrope component is less thanthat in typical mantle-derived garnet.Plagioclase is contained in olivine-hosted melt inclusions

in the olivine basalt fromTakada Ohdake (5072004) and apyroxene-bearing basalt (HM0701D) from Nishimori.The laths of plagioclase are 15^20 mm long. Aluminousclinopyroxene is intimately associated with the plagioclasecrystals in the melt inclusions. The plagioclase is less

calcic (An60^65) than the phenocryst cores (An90^95), andmore sodic than the phenocryst rims (An70).Al^Cr spinel occurs abundantly in the melt inclusions in

olivine in MgO-rich samples, but it does not occur asphenocryst or as a groundmass mineral. The spinel in theMgO-rich samples is chromium-rich, ranging broadlyfrom 0�28 to 0�65 in Cr/(AlþCr). The CrÂl spinelis also rich in Fe2þ relative to Mg, with a 0�45^0�55Fe2þ/(MgþFe2þ) ratio, which is significantly higher thanthat of mantle-derived spinel.The spinel crystals often con-tain round melt inclusions that contain aluminous clino-pyroxene and rare garnet.

WHOLE -ROCK GEOCHEMISTRYMost of the sampled rocks are low-K basalts and basalticandesites; minor medium-K andesite occurs in Nishimori(Table 1 and Fig. 3). Wide ranges in MgO (4^10%), Ni(520^130 ppm), and Cr (30^260 ppm) contents confirmextensive differentiation (Fig. 4). Some samples are high inMgO (8^10%) and low in FeO/MgO (1�0^1�3); these arethe most primitive basalts in the volcanic arc. They are,however, even poorer in Ni and Cr than a magnesianandesite from a nearby cone (Ebisumori), which is not aprimitive magma but the product of mixing between

Table 2: Representative compositions of phenocryst and groundmass minerals fromTakada Ohdake and Nishimori

Mineral: Plagioclase Olivine Orthopyroxene Clinopyroxene

Cone: Nisihimori Nisihimori Nisihimori Takada Nishimori Nishimori Takada Nishimori Nishimori Takada Nishimori Nishimori Takada Takada

Ohdake Ohdake Ohdake Ohdake Ohdake

Sample no.: Nishi-s3 Nishi-s3 Nishi-s3 5072004 Nishi-s3 toak1 5072004 Nishi-s3 Iwht3 5072103N Nishi-s3 Iwht3 5072103N 5972004

Core Rim Gmass Core Gmass

SiO2 45�55 51�46 50�99 44�7 39�9 34�93 38�33 55�21 53�04 53�28 52�55 52�54 52�04 52�69

TiO2 0�08 0�25 0�23 0�18 0�38 0�47 0�22

Al2O3 34�31 29�36 28�11 33�13 0�09 1�74 1�06 1�18 2�47 1�26 2�29 1�04

Cr2O3 0�33 0�13 0�06 0�47 0�10 0�04

FeOt 0�72 0�98 1�31 0�52 17�8 42�76 21�39 11�55 23�23 19�17 4�94 11�10 10�48 16�59

MnO 0�24 0�67 0�37 0�27 0�66 0�38 0�16 0�35 0�39 0�32

MgO 42�4 21�88 39�18 27�89 20�15 23�56 17�32 13�85 14�86 20�99

CaO 19�46 14�57 13�31 18�46 0�15 0�11 0�13 2�27 2�18 2�31 21�04 20�88 19�87 5�53

Na2O 0�67 3�48 3�55 0�73 0�11 0�15 0�15

K2O 0�02 0�14 0�30 0�03

NiO 0�06

An 94�1 69�3 66�3 93�1

Fo 80�7 47�3 76�5

En 77�5 58�0 65�4 49�2 39�5 42�4 60�3

Wo 4�5 4�5 4�6 42�9 42�8 40�1 26�8

Gmass, groundmass.


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primitive basalt, evolved basalt and evolved andesitemagmas (Ohba et al., 2007). All the basaltic rocks werederived from variably differentiated parental magmas.Harker variation diagrams (Fig. 4) show that MgO, Cr,

and Ni contents in rocks from Takada Ohdake decreasewith increasing SiO2 content with concave-down trends.The most magnesian samples in the Nishimori datasetdiffer from those of Takada Ohdake in terms of SiO2 varia-tions. At Nishimori the silica content decreases slightlyfrom 52% to 50�5% concomitantly with a decrease in theNi and Cr contents, then it increases. Rocks fromTakadaOhdake display inflected trends with an early stage peakin CaO and a middle-stage peak in Al2O3. Rocks fromNishimori exhibit similar enrichments in Al2O3 and CaO,but the peaks are obscure on the SiO2 variation diagramsbecause of the back-and-forth behaviour of SiO2, whichcreates a monotonic trend in CaO and a scattered trendin Al2O3.On an AFM (Na2OþK2O^total iron^MgO) diagram,

the differentiation trends of both cones are typical of a tho-leiitic differentiation trend, which is subparallel to theFeO^MgO axis in the early stages of differentiation

(Fig. 5a). A few evolved samples from Nishimori plotslightly towards the total alkali apex, showing an inflectionin the trend of Fe-enrichment. The tholeiitic trends andlow-K characteristics are consistent with the classicisland-arc low-K tholeiite trend reported from northernHonshu (Miyashiro, 1974; Kimura et al., 2002). Increasesin FeO/MgO relative to SiO2 in the early stages of differen-tiation are also typical of a tholeiitic differentiation trend(Fig. 5b). These increases in FeO/MgO on the AFM andMiyashiro diagrams are considered a characteristic ofisland-arc tholeiitic rock series (Miyashiro, 1974); however,iron-enrichment is not observed in the differentiationtrend of Takada Ohdake on the iron^silica variation dia-gram (Fig. 4).Variations in the phenocryst assemblage are roughly cor-

related with bulk-rock compatible element contents atTakada Ohdake (Table 1, Fig. 4). Clinopyroxene-freebasalt lavas are rich in MgO (6�6^10�0wt %), Cr (100^200 ppm), and Ni (40^130 ppm), and low in FeOt/MgO(1�1^1�6). These rocks contain abundant olivine pheno-crysts�minor orthopyroxene. Clinopyroxene-bearingbasaltic rocks are poorer in MgO (4�5^6�8wt %),

Table 3: Representative compositions of minerals in melt inclusions fromTakada Ohdake and Nishimori

Mineral: Clinopyroxene Orthopyroxene Plagioclase Garnet Spinel

Cone: Nishimori Takada Nishimori Nishimori Takada Takada Nishimori Takada

Ohdake Ohdake Ohdake Ohdake

Sample no.: Nishi-s3 HM0701D 5072004 5072004 Nishi-s3 HM0701D 5072004 5072004 5072004 Nishi-s3 5072004

SiO2 43�51 48�88 43�03 42�50 50�64 52�51 55�49 37�14 38�48

TiO2 2�34 2�17 1�49 1�92 0�33 1�06 2�17 1�58 0�88

Al2O3 10�95 7�54 14�00 11�82 6�38 29�79 25�41 21�76 20�49 26�82 32�00

Cr2O3 0�05 0�5 — 0�12 0�10 0�01 0�59 22�22 17�26

FeOt 10�08 8�14 12�07 10�13 16�44 0�67 1�89 20�32 10�64 33�83 34�38

MnO 0�24 0�16 0�33 0�08 0�34 0�19 0�24 0�26 0�18

MgO 10�8 16�34 8�45 8�98 23�29 4�20 7�89 10�1 12�43

CaO 20�05 15�84 19�18 21�71 1�65 12�83 11�63 13�28 16�16 0�02 0�02

Na2O 0�32 0�27 0�22 0�31 — 3�86 4�16

K2O 0�07 0�03

NiO 0�11 0�20

V2O5 0�43 0�42

An 64�5 60�6

En 47�1 58�3 41�2 43�3 26�2

Wo 32�5 25�4 26�3 34�5 4�9

Almandine 45�2 23�3

Grossular 37�8 43�5

Pyrope 16�6 30�8

Cr/(Cr þ Al) 0�36 0�27


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Cr (40^90 ppm), and Ni (20 ^40 ppm) and high in FeOt/MgO (1�5^2�0); the least magnesian samples (c. 4 wt %MgO) are olivine-free rocks. At Nishimori, the correlationbetween chemical composition and mafic phenocrystassemblage is slightly more obscure because no clinopyrox-ene-free samples exist except for a pyroxene-free olivinebasalt sample (toak1). That sample is exceptionally poorin MgO, Ni, and Cr. Olivine-rich samples are richerin MgO (6^8wt % MgO), Cr (100^250 ppm), and Ni(30^70 ppm) than olivine-poor and olivine-free samples(4^6% MgO, 20^80 ppm Cr, and 10^30 ppm Ni).REE abundances are almost 10 times the chondritic

values (Fig. 6), which is typical of island arc tholeiiticbasalts (Wilson,1989). The lowest REE contents are exhib-ited by high-MgO samples from both cones. All theNishimori samples are slightly enriched in light REE(LREE) relative to heavy REE (HREE); chondrite-nor-malized (La/Sm) is 1�1^1�8 and (La/Lu) is 1�4^2�3. Onthe other hand, the REE patterns of the Takada Ohdakesamples are mostly flat. A europium anomaly is weak orabsent in both cones.

DISCUSS IONIntensive parametersMagma temperatures were estimated using pyroxenegeothermometry. The pyroxene geothermometer ofLindsley (1983) was used for groundmass pigeonite and

augite in the augite phenocryst-free high-MgO basaltsample (5072004) from Takada Ohdake, resulting in tem-peratures of 1100^12008C. Presumably, the magma wouldhave been hotter than the maximum pyroxene tempera-ture of 12008C at the time of eruption. Groundmass pyrox-enes crystallized during the final cooling stage at thesurface. The two-pyroxene thermometer of Brey & Ko« hler(1990) was applied to coexisting phenocryst pyroxenes inthe other samples. The pyroxene temperature of an inter-mediate-MgO sample from Takada Ohdake (05072106N,6�5% MgO) is 11208C, and that of an evolved sample(5072103N, c. 5% MgO) is 10908C, indicating that magmatemperature decreases concomitantly with decreasingMgO or increasing degree of differentiation. The two-pyroxene temperature similarly decreases with MgOcontent at Nishimori. The two-pyroxene temperaturesare 1180^11908C for the MgO-rich sample (Nishi-s3),1120^11608C for intermediate-MgO samples (HM0701Dand HM703A, 6�9% MgO; asn6, 5�2% MgO), and9908C for an evolved MgO-poor sample (iwht3,4�3% MgO).Water contents were estimated from the compositions

of plagioclase and whole-rocks using formulations involv-ing relations between pressure, temperature, water con-tent, and plagioclase^melt partitioning (Putirka, 2005).Two-pyroxene temperatures were used to constrain teme-prature. A magma temperature of 12008C was assumedfor the clinopyroxene-free basalt from Takada Ohdake.

Fig. 2. Back-scattered electron images showing occurrences of olivine-hosted crystal-bearing melt inclusions in 5072004. (a) A tiny garnetcrystal occurs in a large aluminous clinopyroxene (cpx) contained in the inclusion next to the clinopyroxene. Glass is mostly devitrified, butsome remains. (b) Clinopyroxene and spinel in a melt inclusion. (c) Plagioclase crystal in a melt inclusion. (d) Garnet crystal in the double-capsule inclusion.The inclusion hosted by spinel is contained in a melt inclusion hosted by olivine. In the inclusion, the garnet crystal aggregateswith aluminous clinopyroxene.


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Whole-rock compositions were assumed to representthe melt compositions equilibrated with the cores of theplagioclase phenocrysts. Estimated water contents are3�1^3�2wt %, uniformly, for the high-Mg samples(5072004 and Nishi-s3), corresponding to a water satura-tion pressure of 7^8 kbar. Intermediate-MgO samples(5�2^6�9% MgO) are slightly richer in H2O (3�6^5�0%)compared with the high-MgO samples, and estimatedH2O contents increase progressively with decreasingMgO. Water contents of 6�5^7�2% were estimated for themost evolved samples (4�2^5�0% MgO). A water contentof c. 2�5% was estimated from the composition of plagio-clase in melt inclusion hosted by olivine in the high-MgObasalt sample, assuming an equilibrium temperature of12008C. Lower water contents (1�0^1�8wt %) are esti-mated when higher magma temperatures (1250^13008C)are assumed for the plagioclase in the melt inclusions.

Equations presented by Putirka (2005) are also useful inestimating the pressures of crystallization, but they includesignificant uncertainties (1^2 kbar or more). Estimatedpressures are 3^5 kbar for most samples, 7 kbar for theprimitive sample (5072004), and 50�5 kbar for the mostevolved sample (iwht3). Pressure estimation based on thesodic plagioclase in melt inclusions was 9^13 kbar, whichis distinctly higher than the phenocryst crystallizationpressures.

Origin of phenocrystsThe order of crystallization, determined using traditionalmicroscopic observation of phenocryst assemblages,involves the early crystallization of orthopyroxene prior toclinopyroxene, as indicated typically by the phenocrystassemblage olivineþ orthopyroxeneþplagioclase in amagnesian basalt. The same crystallization order has beenrecognized for some volcanoes along the volcanic front ofnorthern Honshu by Aoki (1961), who described thecommon occurrence of low-Ca pyroxene during earlystage differentiation of primitive tholeiitic magmas. Inthese volcanoes, the variation in the phenocryst assem-blage is apparently related systematically to bulk-rockchemical variations, which could lead to the conclusionthat the chemical variations resulted from phenocrystremoval. However, quantitative modelling of samplesfrom some basaltic volcanoes has rejected phenocrystremoval as the main process producing the chemical varia-tions (e.g. Ishikawa et al., 1984; Nakagawa, 1985; Ban, 1991).The observed chemical variations within basalts fromIwate and Akita Komagatake require early stage removalof Ca-rich pyroxene from the parental magmas, whereaspetrographic determinations of the order of crystallizationindicate the later crystallization of clinopyroxene(Ishikawa et al., 1984; Nakagawa, 1986).Similar to Iwate, the removal of observed phenocrysts is

not the dominant process in producing the compositionalvariation at either Takada Ohdake or Nishimori. AtNishimori, CaO/Al2O3 decreases monotonously and stee-ply with decreasing MgO (Fig. 7), which implies dominantfractionation of clinopyroxene. This is inconsistent withthe petrographic data as the augite phenocryst content islow, or even absent, and the major phenocrysts are plagio-clase, olivine, and orthopyroxene. The geochemical trendat Takada Ohdake shows a peak in CaO/Al2O3 in theearly stages (Fig. 7), but neither the early stage increasenor the later stage decrease can be explained by removalof plagioclase, olivine, and orthopyroxene phenocrysts.Augite occurs only in the more differentiated samples.Therefore, we conclude that phenocryst removal is insuffi-cient to account for the variation trend of Takada Ohdake.Because the observed phenocryst minerals are not

the fractionating phases in the evolving magma bodies,the fractionated phases are considered to be cryptic,possibly extinct, or only present as inclusions or cores

Fig. 3. (a) Total alkali vs silica (TAS) diagram (after Le Baset al., 1986). The analysed samples are mostly basalts and basalticandesites. One sample from Nishimori is classified as an andesite.(b) Subdivision of sub-alkaline rocks on a K2O vs SiO2 diagram(Le Maitre, 2002). Most samples plot in the low-K field, but twosilica-rich samples plot in the medium-K field.


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of phenocrysts. Visible phenocrysts are mostly later-crystallized minerals. Crystallization in a shallow magmachamber is a probable petrogenetic process, but we prefera simpler model of decompression-dehydration crystalliza-tion during ascent (Brophy et al., 1999; Kuritani, 1999;Blundy & Cashman, 2001, 2005; Annen et al., 2006),which is the same mechanism as microlite crystallization(e.g. Couch et al., 2003). The liquidus temperature risesand the consequent supersaturation promotes crystalgrowth when H2O vapour is released from a water-saturated magma during decompression. An initialhydrous magma condition is necessary for this model.The high water contents estimated for these volcanoes areconsistent with such a dehydration model.

Although a large fraction of the phenocryst populationis considered to have crystallized at pressures less than10 kbar where garnet is unstable, possibly attributable todecompression dehydration, some phenocrysts might haveoriginated from greater depths because they contain meltinclusions of minerals that are stable at high pressure.Melt inclusions in the olivine basalt (5072004) fromTakada Ohdake contain garnet, aluminous clinopyroxene,and plagioclase, seemingly implying garnet granulite-facies conditions (10^25 kbar). As discussed in the preced-ing section, plagioclase in these melt inclusions crystallizedat 9^13 kbar, suggesting a deeper location than thatfor phenocryst plagioclase. Therefore, although pheno-crysts form mostly at shallow depths (510 kbar), some

Fig. 4. Selected major element variations vs SiO2 forTakada Ohdake and Nishimori. Iron-enrichment, a characteristic of a tholeiitic differenti-ation trend, is not observed within the Takada Ohdake samples on the FeOt^SiO2 diagram, but they are regarded as a tholeiitic rock seriesbased on the variation trend on the AFM and Miyashiro diagrams (Fig. 5).


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phenocrysts trapped and crystallized melt inclusions atpressures higher than c. 10 kbar. The phenocrysts might bederived from a deep magma chamber in which crystalliza-tion differentiation was occurring.

Trace element modellingWe evaluated quantitatively whether crystallization differ-entiation is responsible for the range of chemical variationin the following manner. Assuming perfect (Rayleigh)fractional crystallization, we adopt the equation proposedbyAlle' gre & Minster (1978):

C ¼ C0FD�1: ð1Þ

In the equation, C signifies the concentration of an elementin the melt, C0 is the initial concentration, F represents themelt fraction remaining and D is the bulk mineral^melt

partition coefficient. The equation can be expressed for aselected reference element h as

Ch ¼ Ch0F

Dh�1: ð2Þ

In this equation, Ch is the concentration of the referenceelement in the melt, Ch

0 is the initial concentration ofthe element, and Dh is the bulk partition coefficient of thereference element h. By eliminating F from equations (1)and (2), the relation between these two elements can beexpressed as

ch

c¼

chð1�DÞ=ð1�Dh Þ

0

c0ChðD�D

h Þ=ð1�Dh Þ

: ð3Þ

This is a power function with variables of Ch and Ch/C. Thevariation trend can be fitted by the power function (3) ona diagram of Ch vs Ch/C if perfect fractional crystallizationis responsible for the chemical evolution. The exponent ofthe fitted equation is a function of the partition coefficients,(D ^ Dh)/(1 ^ Dh).The exponent approximates to the parti-tion coefficient D for many elements because Dh is approx-imately zero when a highly incompatible element ischosen as the reference element (Alle' gre & Minster, 1978).Potential fractionating minerals can be inferred by com-paring the fitted values of (D ^ Dh)/(1 ^ Dh) (expressed asD0 hereafter) and reported partition coefficients betweenminerals and melt. One or more minerals must have ahigher D0 than the fitting estimation or bulk D0 if the frac-tionated solid phase involves more than one mineral.The estimation is carried out for as many elements as pos-sible to narrow down the potential fractionating minerals.For this purpose, high but not extraordinary values ofthe partition coefficients were selected for comparisonfrom previous studies, except for clinopyroxene. For clino-pyroxene, REE partition coefficients were calculatedusing the method described by Wood & Blundy (1997).In our modelling, we chose La and Rb as reference ele-ments. Hereinafter, (D ^ DLa)/(1 ^ DLa) and (D ^ DRb)/(1 ^ DRb) are expressed respectively as D0Rb and D0La.Calculated bulk D0 values for selected trace elements are

given in Table 4. We excluded two andesite samples (asn6and iwht1), which are later stage differentiation productsafter the Fe-enrichment peak on the AFM diagram(Fig. 5), to reduce the complexity of the phase changesduring progressive differentiation.Therefore, our trace ele-ment modelling corresponds to the early stages of tholeiiticbasalt differentiation that involves an increase of FeO/MgO with fewer increases in SiO2 or Na2OþK2O.Chemical trends on diagrams of Rb vs Rb/C and La vsLa/C have positive gradients for many tested elements(Fig. 8), implying that the distribution coefficients of theseelements are larger than those of Rb and La. The fittedbulk D0Rb values are 0�7^1�1 for Sr, Sc, Y, Zr, Hf, Nb,Ta, and the REE. These values are considerably higherthan typical D0Rb for olivine, spinel, orthopyroxene,

Fig. 5. (a) AFM (Na2OþK2O^FeO^MgO) diagram.The variationtrends of both cones are typical of the tholeiite rock series, exhibitingiron-enrichment relative to MgO. An iron-enrichment peak is distinctat Nishimori. (b) Miyashiro (FeOt/MgO vs SiO2) diagram. Arrowsare inferred differentiation trends; dashed line: Takada Ohdake,dotted line: Nishimori. Slopes of the trends in the early stages ofdifferentiation are steeper than that of boundary between TH(tholeiitic) and CA (calc-alkaline) defined by Miyashiro (1974).


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Fig. 6. Chondrite-normalized REE patterns for (a) Takada Ohdake and (b) Nishimori [normalizing values fromTaylor & McLennan (1985)].

Fig. 7. Chemical variation trends shown on MgO (wt %) vs CaO/Al2O3 diagrams. Gray arrows indicate the variation trends of TakadaOhdake and Nishimori. Expected differentiation trends by removal of minerals are indicated by black arrows. Dashed arrows indicate fractiona-tion of the phenocryst minerals observed in MgO-rich samples (olivine, anorthitic plagioclase, and orthopyroxene). The continuous-linearrows indicate fractionation of aluminous clinopyroxene and sodic plagioclase.

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and clinopyroxene. Modelling using La instead of Rbresulted in appreciably flatter slopes for most elements,implying that DLa is considerably greater than DRb,although previous studies of partition coefficients haveshown both elements to be highly incompatible. The D0Rb

value calculated for La is c. 0�8.As with the major elements, the trace element modelling

is inconsistent with the removal of the observed phenocrystphases as the major differentiation process (Fig. 9a). TheD0 values of many elements cannot be accounted for byfractionation of the major phenocryst minerals (olivine,plagioclase, orthopyroxene, and spinel). Anomalous D0

values for Sr and Eu are attributed to the fractionation ofplagioclase. The high D0Rb values for Nb, Hf, Zr, and Taare considered to result from spinel fractionation becausethese elements are compatible in spinel (Nielsen et al.,1993). Rutile and ilmenite also have high partition coeffi-cients for these elements, but at magmatic temperaturescrystallization of these minerals is less likely. One or moreREE-compatible cryptic minerals must be included inthe fractionated mineral assemblage along with plagio-clase and spinel to account for the D0 values of the REE.

Calcic pyroxene, amphibole and garnet are suitablecandidates.To compare the D0 of clinopyroxene with the estimated

bulk D0, the method of Wood & Blundy (1997) was appliedto estimate partition coefficients between a basalticmagma with the composition of 5072004 and clinopyrox-ene with two compositions: (1) phenocryst augite; (2) alu-minous clinopyroxene in the melt inclusions. The profileof D0Lacpx for the REE (Fig. 9b) shows low values forthe LREE and higher relatively constant values for themiddle REE (MREE) to HREE. The profile closelyresembles the bulk D0La for both Takada Ohdake andNishimori, except for the Eu spike, which is considered tobe a consequence of plagioclase fractionation. Further-more, the D0La values of clinopyroxene are all higher thanthe bulk D0La. Therefore, it is highly probable that clino-pyroxene is the major fractionating phase. The profiles ofD0Rb

cpx for the REE resemble those of D0Lacpx, but theslope of the LREE values is steeper than the estimatedbulk D0 profiles (Fig. 9c). Nevertheless, the overall shapesof the bulk D0Rb and D0Rb clinopyroxene profiles are simi-lar. Green et al. (2000) pointed out that an increase in theAlIVcontent of clinopyroxene enhances the partition coef-ficients of the REE. The aluminous pyroxene has suffi-ciently high D0Rb to account for the bulk D0Rb, whereasthe D0Rb values for augite may be too low. The mineralchemistry evidence for the crystallization of AlIV-rich clin-opyroxene is consistent with this trace element modelling.REE partition coefficients for amphibole show a similar

profile (not shown), but with higher absolute values thanthose of clinopyroxene (Fig. 9c). The melt^mineral parti-tion coefficients of both clinopyroxene and amphibolevary considerably according to melt and mineral composi-tions. Therefore, which mineral is the dominant fractionat-ing phase cannot be determined from this trace elementmodelling. However, we can exclude the possibility ofamphibole fractionation in the petrogenesis of these basal-tic volcanoes. The estimated two-pyroxene temperaturesare higher than 11008C and are typically 12008C. At suchtemperatures, amphibole is not stable in basaltic magmasat any pressure, even if the magmas are saturated withH2O. Maximum crystallization temperatures of amphi-bole are typically 900^10508C, according to meltingexperiments on basaltic rocks under hydrous conditions(Allen & Boettcher, 1983; Sen & Dunn, 1994; Rapp &Watson, 1995; Nakajima & Arima, 1998). Therefore, weconclude that the fractionated phases from the cones didnot include amphibole.The occurrence of garnet occurrence in the melt inclu-

sion assemblages implies that the magmas commencedcrystallization at pressures at which garnet is stable(greater than c. 10 kbar) and crystallization differentiationmight therefore have involved garnet fractionation.However, the D0 REE values of garnet differ considerably

Table 4: Calculated D’ values

D0Rb D0La

Nishimori Takada Ohdake Nishimori Takada Ohdake

Sr 1�14 1�10 0�88 0�95

Sc 0�94 1�09 0�94 1�07

Y 1�02 0�93 0�36 0�34

La 0�84 0�80 — —

Ce 0�84 0�81 0�05 0�03

Pr 0�94 0�85 0�11 0�10

Nd 1�03 0�90 0�23 0�23

Sm 1�04 0�93 0�32 0�31

Eu 1�06 1�10 0�52 0�74

Gd 1�04 0�91 0�37 0�29

Tb 1�04 0�92 0�41 0�35

Dy 1�04 0�92 0�41 0�32

Ho 1�05 0�91 0�36 0�34

Er 1�03 0�89 0�41 0�29

Tm 1�02 0�86 0�42 0�28

Yb 1�04 0�85 0�44 0�27

Lu 1�04 0�86 0�45 0�28

Zr 0�96 0�72 �0�15 �0�16

Hf 1�03 0�72 �0�01 �0�18

Nb 0�91 1�04 0�24 0�22

Ta 1�09 1�14 0�42 0�2

Th 0�92 0�13 �0�22 �1�3


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Fig. 8. Examples of trace element modelling to determine D0 using the Ch vs C/Ch diagrams (Ch is the concentration of the reference element).La and Rb are used as reference elements. The definition and determination method for D0 are given in the text. The data define trends withpositive gradients for many incompatible elements; accordingly, high D0 values are estimated.


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from those estimated for the sampled rocks (Fig. 9b),suggesting that garnet is not a major fractionatingphase. The D0La values of garnet increase progressivelyfrom the LREE to the HREE, resulting in much higherD0La HREE/D0

LaLREE, whereas the bulk D0 profiles have

plateaux from the MREE to HREE. It is probable there-fore that garnet was not part of the fractionating assem-blage. Even if garnet were fractionated, it would accountfor a very small proportion of the entire fractionatingassemblage and would therefore have minimal effect onD0 values.Our results indicate that the trace element characteristics

of the magmas can be accounted for by fractional

crystallization involving plagioclase, aluminous calcic pyr-oxene, and spinel, which does not discount the involvementof minerals with low partition coefficients (e.g. olivine andorthopyroxene) fromthe fractionated solidphase.Their par-ticipation has little effect on thebulk partition coefficients.In the next section, the equilibrium phase relations of

the fractionating minerals are discussed using MELTSmodelling (Ghiorso & Sack, 1995) along with the depth ofcrystallization differentiation.

Depth of crystallization differentiationThe first step in estimating the depth of fractional crystalli-zation of the magmas necessitates discussion of the mineral

Fig. 9. Profiles of estimated bulk D0 values for Takada Ohdake (circles) and Nishimori (squares) compared with D0 values for potentiallyfractionated minerals from the literature (except cpx). Data sources of partition coefficients: Green (1994); Taura et al. (1998); Green et al.(2000); Adam & Green (2006); Shaw (2006); Be¤ dard (2007). Partition coefficients for clinopyroxene were estimated using the methoddescribed byWood & Blundy (1997). (a) Comparison between bulk D0Rb and those of phenocryst minerals in magnesian samples (olivine, ortho-pyroxene, plagioclase, and spinel). (b) Bulk D0La for REE and those of clinopyroxene and garnet. (c) Bulk D0Rb for REE and those ofclinopyroxene.


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assemblage, which depends on pressure. The results of thetrace element modelling have demonstrated that therange of bulk-rock chemical variation can be accountedfor by fractional crystallization involving plagioclase, alu-minous calcic pyroxene and spinel. Garnet fractionationis not required by the modelling; however, the occurrenceof this mineral in the melt-inclusions is of significance inindicating a stage of high-pressure crystallization. Garnetcrystals effectively occur in ‘double capsules’ of which theinner parts comprise spinel or clinopyroxene suspended inmelt that was trapped by olivine phenocrysts. Because thehost olivine composition is not primitive, the melt musthave already differentiated somewhat when it was trapped.These melt inclusions are rich in crystals and their intersti-tial melt are more evolved in composition (andesite withc. 60% SiO2) than the whole-rocks. This implies thatthe daughter minerals were partly crystallized from theentrapped melt during the final cooling stage at the sur-face. Assuming that the melt inclusions remained over-pressurized, garnet crystallization in the melt inclusionscould have taken place at the surface. Nevertheless, thepressure of melt inclusion entrapment must be in the pres-sure range of garnet stability in mafic magmas.The depth of crystallization differentiation can be

approximated from the stabilities of minerals by compari-son with basalt melting experiments. The lower pressurelimit corresponds to the limit of the garnetþmelt stability,which is 8^24 kbar (Allen & Boettcher, 1983; Baker &Eggler, 1983; Elthon & Scarfe, 1984; Rapp & Watson, 1995;Liu et al., 1996; Nakajima & Arima, 1998), depending onwhole-rock composition, especially volatile components.The upper pressure limit corresponds to the limit of plagi-oclase stability at basaltic magma temperatures, becauseplagioclase is a fractionated phase. Plagioclase stabilitystrongly depends on magma H2O content, which causesa decrease in the temperature stability of plagioclase.In some hydrous experiments, plagioclase is not stable atbasalt magma temperatures (1100^13008C) at pressuresgreater than 8 kbar (e.g. Nakajima & Arima, 1998). Onthe other hand, anhydrous experiments exhibit broadstability fields for plagioclase from atmospheric pres-sure to 20 kbar (e.g. Baker & Eggler, 1983) at around1100^12008C. This leads us to the conclusion that theparental magmas beneath the volcanoes were probablyanhydrous or near-anhydrous and that the magma crystal-lization pressure was roughly 8^20 kbar.Trace element modelling must be consistent with major

element variations that are controlled by the fractionatingphase assemblage. The phase relations are intricatelyrelated with the bulk composition of the magma, its watercontent, temperature, and pressure. To constrain suitableconditions to account for the major element variations,isobaric fractional crystallization was modelled usingMELTS (Ghiorso & Sack, 1995), for a range of pressures,

oxygen fugacities and water contents. Simulations werecarried out under the following constraints: (1) the fractio-nating mineral assemblage includes aluminous clinopyrox-ene, plagioclase and spinel; (2) minor garnet crystallizesfrom the melt at any temperature; (3) the models shouldreproduce the early stage peaks in Al2O3, middle-stagepeaks in CaO and early stage depletion and middle-stageplateaux in total iron.The MELTS simulation results are presented as a series

of Harker diagrams in Fig. 10. No model result was a per-fect match, but the model curves shown in the figuresclosely approximate the data for each volcano. For TakadaOhdake, the best matching trends were simulated at10 kbar, QFM bufferþ 3�2 log units, and anhydrous condi-tions. Higher pressures and water contents reduce the tem-perature of plagioclase crystallization, thereby delayingthe Al2O3 peak. Under these conditions, garnet crystal-lizes from the differentiated melt at c. 10508C and the resid-ual melt becomes andesitic, similar to the observation.At lower pressures, garnet does not crystallize at any pres-sure. Variation in the silica content is influenced stronglyby oxygen fugacity; the observed silica variation isapproximated at QFMþ 3�5. In a similar manner, thebest-model condition of 10 kbar, anhydrous, and QFMbuffer ^ 1�0 log units was determined for Nishimori. Thesimulation at these conditions accounts for the early tomiddle stages of the observed trends, but a significantdiscrepancy exists for FeO during late-stage differentiation.Assuming that a change in oxygen fugacity from QFM ^1 to QFMþ 3 occurred in the middle stage of differentia-tion, the simulation can account for the decreasing trendof iron. In these preferred models, the fractionated miner-als are aluminous calcic-clinopyroxene, plagioclase, andspinel. Orthopyroxene is a liquidus mineral but is stableonly at near-liquidus temperatures. The simulated mineralassemblage is also consistent both with the trace elementmodelling and the petrography of the melt inclusions. Theassumption of simple fractional crystallization might bean oversimplification and more complex differentiationmodels might approximate the observed trends muchbetter, but we believe that the essence of the result is sub-stantiated by the consistency among the trace elementmodelling, petrographic observations and the MELTSmodelling.The results of this study generate a new problem related

to the behaviour of H2O. The phase relation simulated byMELTS requires anhydrous conditions, as predicted byseveral experimental studies. For example, Hamada &Fujii (2008) demonstrated that tholeiitic trends can be pro-duced experimentally with less than 2wt % H2O. Theplagioclase composition in the melt inclusions (An60) isconsistent with such an anhydrous condition. On theother hand, hydrous conditions were estimated fromequilibria between phenocryst plagioclase (An88^95) and


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the magma. Such water-rich conditions for the petrogen-esis of arc tholeiite basalts have been recognized on thebasis of experimental studies and the petrography of melt-inclusions (Takagi et al., 2005; Hamada & Fujii, 2007).We do not consider this discrepancy a contradiction, butrather a result of changes in the water content of themagmas with progressive differentiation. The magmaswere initially anhydrous; they then became hydrous aftercrystallization differentiation. The inferred water contentsof the later stage magmas are too high to be accounted forby fractional crystallization of anhydrous phases.We there-fore conclude that these magmas must have gained H2Ofrom the surrounding country rocks. The water couldhave been released from local amphibolite wall-rocks thatwere metamorphosed to granulite facies when heated by

the magma, or from older solidifying magma bodies (seeDavidson et al., 2007).The estimated pressure of 10 kbar corresponds to a litho-

static depth of c. 34 km, which is similar to that of theMoho in the region (35^36 km). Emplacement of basaltmagmas at the Moho is a classical concept often calledunderplating. Results of recent studies (Annen & Sparks,2002; Annen et al., 2006) show a link between underplatingand the genesis of intermediate and silicic magmas,demonstrating that basalt intrusion around the Mohogenerates a hot crustal zone where silicic and intermediatemagmas are produced by either crustal melting or crystal-lization differentiation. This model is consistent withthe observation that silicic and intermediate magmas aremajor constituents of the volcanoes surrounding Takada

Fig. 10. Major element fractional crystallization modelling with the MELTS simulation shown on Harker diagrams for Takada Ohdake(circles) and Nishimori (squares). Lines show some models that are well matched to the observed trends; continuous lines: Takada Ohdake;dashed lines: Nishimori. The model constraints are described in the text. Symbols are subdivided according to the mafic mineral assemblageas shown in the legend. Fractionating mineral assemblages modelled with MELTS are different from the visible phenocryst assemblages.


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Ohdake and Nishimori. Such thermal interaction betweenmagma and crustal rocks might involve migration ofwater into the magma.

CONCLUSIONS

(1) Melt inclusions hosted by olivine in basalts fromTakada Ohdake and Nishimori contain aluminousclinopyroxene, sodic plagioclase, spinel, and raregarnet.

(2) The crystallization sequence of mafic minerals(olivine ! orthopyroxene ! clinopyroxene) deter-mined by petrographic observations of phenocrystminerals is not consistent with the chemical variation.The phenocryst minerals are therefore probably notformed in the magma chamber in which crystalliza-tion differentiation occurred. Possibly, they crystal-lized by decompression-induced degassing.

(3) High distribution coefficients between the melt andthe fractionating solid phase were estimated for manytrace elements.The data can be accounted for by frac-tionation of aluminous clinopyroxene, spinel, and pla-gioclase. Garnet is not involved in the fractionationassemblage; rather, it crystallized during the finalcooling stage at the surface. High pressure (10 kbar)was maintained in the melt inclusions until eruption.

(4) The depth of crystallization differentiation is esti-mated as c. 34 km, which is approximately the depthof the Moho in the region (35^36 km).

(5) The parental basalt magmas were initially anhydrous.They became hydrous before phenocryst crystalliza-tion at shallow depths.

ACKNOWLEDGEMENTSThis work was partly carried out as part of their Mastersthesis studies byY. Kimura and K. Matsuoka and as a doc-toral thesis study by T. Ohba at the Graduate School ofScience, Tohoku University. The authors acknowledgeY. Oyama, Y. Ito, and H. Kawanobe for their technicalsupport; K. Aoki, K. Onuma, S. Kanisawa, T. Yoshida,T. Hasenaka, and N. Tsuchiya for constructive discussions;and Y. Kanahara for his support during the field survey.Discussions with R. S. J. Sparks, S. Mahoney, andC. Martin in the field were extremely helpful. We are alsograteful to I. E. Smith for his kind help in improving themanuscript. The first draft of this manuscript was writtenby the senior author (T.O.) during his sabbatical stay atThe University of Auckland. We thank P. R. L. Browne,K. B. Spo« rli, and I. E. Smith of the University ofAuckland, and the 21st Century COE program(Advanced Science and Technology Center for theDynamic Earth) of Tohoku University for supporting hissabbatical. Finally, we would like to express our deepappreciation to Professor Wilson, the Executive Editor

of the journal, for her kind efforts in improving themanuscript.

SUPPLEMENTARY DATASupplementary data for this paper are available at Journalof Petrology online.

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