JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 102, NO. B4, PAGES 8085-8103, APRIL 10, 1997

High B (HIMU)ocean island basalts in southern Polynesia: New evidence for whole mantle scale recycling of subducted oceanic crust

Tetsu Kogiso, 1 Yoshiyuki Tatsumi, 2 Gen Shimoda, 3 and Hans G. Barsczus 4's

Abstract. Major elements, trace elements, and Pb isotopic compositions were determined for ocean island basalts (OIBs) from Polynesia in the southern Pacific in order to document the chemi- cal characteristics of OIB sources and to understand their origin. High [.t (HIMU: [.t = 238U/2ø•pb) basalts, which have distinctly high Pb isotopic ratios, have systematically different compositions from non-HIMU basalts; HIMU basalts are more enriched in Fe203*, MnO, and CaO and more depleted in SiO2, K20, P205, Ni and incompatible trace elements than non-HIMU, except for Nb. Major element characteristics of HIMU basalts suggest that the HIMU source is more fertile, i.e., more enriched in a basaltic component, than non-HIMU sources. This is consistent with the sug- gestion that subducted oceanic crust may contribute to the formation of the HIMU reservoir. Rela- tive depletion of incompatible trace elements in HIMU is consistent with involvement of sedimen- tary components in non-HIMU sources. However, enrichment of Nb relative to other incompatible elements in HIMU cannot be explained by simple addition of the crustal component nor partial melting processes in the upper mantle, implying that lower mantle processes may contribute to the formation of the HIMU source.

Introduction

Numerous models have been proposed to explain the wide diversity of trace element and isotopic compositions of ocean island basalts (OIBs). Four major end-member compositions, depleted mid-oceanic ridge basalt (MORB) Mantle (DMM), EnrichedMantle 1 and 2 (EM1 andEM2), and high • (HIMU = high 238U/2ønpb), are generally invoked to explain the isotopic variations in OIB suites [Zindler and Hart, 1986; Weaver, 1991 ]. Among these four components, the HIMU component is characterized by extremely high 2ø6pb/2ønpb, 2ø7pb/2øapb, and 2øspb/2ønpb, and MORB-like 87Sr/S6Sr and 143Nd/lnaNd. The most typical HIMU basalts (2ø6pb/2ø4pb >20.5)occur in two regions on the Earth, St. Helena in the southern Atlantic Ocean and Cook-Austral Islands in the southern Pacific Ocean. P wave

whole mantle tomography demonstrated that the whole mantle scale upwelling is present beneath these two regions [Fukao et a/., 1994]. This implies that the origin of the HIMU compo- nent may be related to the dynamic processes occurring in the deep mantle.

The extremely radiogenic Pb isotopic ratios of HIMU basalts suggest an enrichment both in U and Th relative to Pb

1 Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Tokyo, Japan.

2 School of Earth Sciences, Faculty of Integrated Human Studies, Kyoto University, Kyoto, Japan.

3 Graduate School of Human and Environmental Studies, Kyoto Uni- versity, Kyoto, Japan.

4 G•ofluides-Bassins-Eau, Institut des Sciences de la Terre, de l'Eau et de l'Espace de Montpellier, Centre National de la Recherche Scienti- fique and Universit6 de Montpellier II, Montpellier, France.

5 Also at Institut Franqais de Recherche Scientifique pour le D6veloppement en Coop6ration (ORSTOM), Montpellier, France.

in the magma source of the order of 109 years ago [Tatsurnoto, 1978; Hauri and Hart, 1993]. These enrichments are not asso- ciated with an increase of Rb/Sr, because HIMU basalts are

characterized by relatively low 87Sr?aSr. Various mechanisms have been proposed for the origin of the HIMU source includ- ing selective extraction of Pb with chalcophile character from the mantle to the core [Volliner, 1977; Vidal andDosso, 1978; All•gre et al., 1980] and metasomatism by CO2-rich fluids [Hart et al., 1986; Nakamura and Tatsurnoto, 1988]. The most widely accepted mechanism would be the involvement of sub- ducted oceanic crust, which has been depleted in Pb relative to U and Th by hydrothermal alteration processes at mid-oceanic ridges and/or during dehydration processes in subduction zones [e.g., Chase, 1981; Hofmann andWhite, 1982; Weaver, 1991; Chauvel et al., 1992]. Experimental studies [Brenan et al., 1995; Kogiso et al., 1997] have demonstrated that U/Pb and Th/Pb ratios of the dehydrated subducted oceanic crust are high enough to account for the observed high Pb isotopic ratios of the HIMU source. If this is the case, then major element com- positions of HIMU magmas are expected to reflect entrainment of such basaltic crustal material in the source. However, previ- ous studies on the origin of OIB have been limited to their iso- tope and trace element characteristics, whereas constraints on the origin of HIMU from major element characteristics have not been systematically investigated.

In this paper, new data of major elements, trace elements, and Pb isotopic ratios are reported on alkalic basalts of HIMU and other non-HIMU lavas from Polynesian islands. The paper proposes a possible mechanism for producing the HIMU source in the deep mantle based on the chemical characteristics (particularly major elements) of HIMU and non-HIMU basalts.

Copyfight 1997 by the American Geophysical Union.

Paper number 96JB03892. 0148-0227/97/96JB-03892509.00

Tectonic and Geological Background

The islands of Polynesia are located in the South Pacific Ocean (Figure 1), where they form five main island chains

8O85

8086 KOGISO ET AL.: HIGH g (HIMU) • ISLAND BASALTS

5øS

10 ø

20 ø

160 ø 150 ø 140 ø 130øW I i i • i i I i i

øø • f4000 ) •"h• Ua Pou \.. 0 o o'o /•.• r Hiva Oa • . ' -)Mani•ki _•.•e •'•- --•• - )Plateau

Bora Bora',

Mangaia • • . er

• Macdonald

30 ø I I I I I I I 160 ø 150 ø 140 ø 130øW

5øS

10 ø

20 ø

30 ø

Figure 1. Location of the studied islands on a map of Polynesia in the southern Pacific Ocean. Contour lines are drawn at 0 and 4000 m.

(Marquesas, Tuamotu, Society, Pitcairn-CJambier, and Cook- Austral), each roughly trending in a northwest-southeast di- rection. The area of Polynesia is characterized by unusually high heat transfer from the deep mantle, which is suggested by (1) shallow ocean floor (" super swell" [McNutt and Fischer, 1987]), (2) lithosphere with unusually thin elastic thickness [MenardandMcNutt, 1984; Calmant and Cazenave, 1987], (3) high geoid residuals [Richards and Hager, 1988], and (4) the location of a large number of hotspots. Further, the whole mantle seismic tomography data [Fukao, 1992] have demon- stratedthe occurrence of a low-velocity anomaly region down to the core-mantle boundary under the Polynesian islands. These observations have lead many authors [Larson, 1991; Coffin and Eldholm, 1994; Fukao et al., 1994] to the specula- tion that a large-scale upwelling of the deep mantle matehal (" superplume" [Larson, 1991 ]) may occur under the Polyne- sian region in the southern Pacific.

The islands of Polynesia are dominantly made up of basaltic rocks (except for Tuamotu Islands, which are coral atolls) with subordinate amounts of differentiated rocks. Volcanic activi-

ties has persisted from about 20 Ma to the present at several submarine volcanoes [Duncan and McDougall, 1976]. Vol- canic rocks of Polynesian islands display a wide variety of iso- topic and trace element compositions; total Sr, Nd, and Pb iso- topic variations within Polynesia are nearly as large as in all OIBs on Earth [e.g., Vidal et al., 1984; Nakamura and Tatsu- moto, 1988; Dupuy et al., 1989].

Samples examined in this study were alkalic basalts from Mangaia, Rurutu, and Tubuai in the Cook-Austral; Bora Bora, Huahine, and Tahiti in the Society; and Ua Pou and Hiva Oa in the Marquesas Archipelago. The Cook-Austral Archipelago

consists of more than 10 volcanic islands and several

seamounts, linearly arrayedin a northwest-southeast direction. There is an active submarine volcano (Macdonald seamount) at the southeastern end of the volcanic chain (Figure 1). Decreas- ing ages of volcanism from 19.5 Ma at Mangaia to 0 Ma at Macdonald seamount indicate the migration of volcanism from northwest to southeast, with a migration rate of 11 cmJyr [Duncan and McDougall, 1976]. However, much younger ages than predicted from the hotspot theory have been reported from some northwestern islands (e.g., Aitutaki, Atiu, Raro- tonga, and especially Rurutu), possibly suggesting the exis- tence of other hotspots along the island chain [Turnerand Jar- rard, 1982; Barsczus et al., 1994]. In Rurutu, two phases of volcanism have been recognized as follows [Duncan and McDougall, 1976]: rocks of the recent series (1.8-0.6 Ma) have slightly lower Pb isotopic ratios than those of the older series (13-12 Ma) which have Pb isotopic ratios typical of HIMU value [Barsczus eta/., 1986]. The ages of the volcanism in Mangaia and Tubuai are 19.5-17 Ma and 12-7 Ma, respec- tively [Duncan and McDougall, 1976; Turner and Jarrard, ]9821.

The Society Archipelago consists of 10 volcanic islands and several seamounts, linearly arrayed in a northwest- southeast direction. The Society hotspot is located in the southeastern end of the archipelago, where some active sub- marine volcanoes are observed[Talanch'erandOkal, 1984]. The ages of the islands decrease from 4.5 Ma at Maupiti to 0.3 Ma at Tahiti, indicating a migration of volcanism with a migra- tion rate of 11 cmJyr [Duncan andMcDougall, 1976; Diraison et al., 1991]. The ages of the volcanism are 3.8-3.2 Ma in Bora Bora, 3.2-2.1 Ma in Huahine, and 1.7-0.3 Ma in Tahiti

KOGISO ET AL.: HIGH g (HIMU) OCEAN ISLAND BASALTS 8087

[Duncan and McDougall, 1976; Diraison et al. , 1991; Duncan et al., 1994].

The Marquesas Archipelago, the northernmost archipelago of Polynesia in the South Pacific, includes 12 volcanic islands and more than eight seamounts. The ages of islands decrease from northwest (6 Ma at Eiao)to southeast (1.3 Ma at Fatu Hiva) [Duncan and McDougall, 1974; Duncan et al., 1986; Brousse el al., 1990], and the ageof 0.4 Ma has been reported for a small seamount 50 km southwest ofFatu Hiva [Desonie el al., 1993]. These observations suggest that the island chain was formed by volcanic activity of a postulated hotspot (Marquesas hotspot), although active volcanism has still not been recognize& The ages of the volcanism are 5.6-1.8 Ma in Ua Pou and 3.2-1.6 Ma in Hiva Oa [Duncan and McDougall, 1976; Duncan el al., 1986; Katao el al., 1988; Brousse et al., 1990].

Analytical Methods

Samples of Rurutu, Tubuai, Mangaia, and Ua Pou were col- lected by the authors (T. K. and H. G. B.), others sample being supplied from H. Morinaga of Himeji Institute of Technology, Japan. From about 500 available samples, 260 relatively fresh rocks were selected for analyses of major and trace elements. Each sample was crushed into small fragments using a hammer. The rock fragments (about 100 g) were crushed to powder (<200 mesh) using a tungsten carbide vessel and shaken in an automatic shaker for 4 to 5 min in order to homogenize the powder. Major elements, Rb, St, Ba, Pb, Th, Nb, Zr, Y, and Ni were analyzed using RIGAKU © 3070 and Symaltics 3550 X ray fluorescence (XRF) spectrometers at Kyoto University by the method described by Goto and Tatsumi [1992, 1994]. Rare earth element (REE) data were obtained for 21 samples by in- ductively coupled plasma (ICP) at Royal Holloway, University of London.

Pb isotopic compositions were measured by a silica gel/phosphate method using a Finnigan © MAT261 single- collector mass spectrometer at Kyoto Sangyo University fol- lowing the methodof ShimodaandNohda [1995]. Each sample for Pb isotope analyses was crushedinto tiny fragments (about 1 mm in diameter). After washing in double-distilled water us- ing an ultrasonic bath, the fragments (150-500 mg) were leached with triple-distilled 2N HC1 for 2 hours at 60øC and then dissolved in ultrapure HF and HC104. Total blank for Pb was typically 50 pg. Measured isotopic compositions were corrected for fractionation of 0.11% per atomic mass unit for 2ø6pb/2ø4pb and 2ø7pb/2ø4pb and 0.13% for 2ø8pb/2ø4pb ratios

based on the replicate analyses of National Institute of Stan- dards and Technology (NIST) 9 81.

tively low MgO contents contain plagioclase as well as oli- vine and/or clinopyroxene phenocrysts. Amphibole phe- nocrysts are present only in evolved rocks. Thus the crystalli- zation sequence of major phenocryst phases during magmatic differentiation was olivine • clinopyroxene • plagioclase • amphibole. Olivine phenocrysts are occasionally surrounded by iddingsite. Clinopyroxene phenocrysts usually show sector and oscillatory zoned structure. Kinked and/or resorbed olivine with chromite inclusions are found in lavas with high MgO contents. Inclusions of chromite and titanomagnetite occur in olivine and clinopyroxene crystals and/or as microphe- nocrysts in some rocks. The groundmass consists mainly of clinopyroxene, olivine, plagioclase, alkali feldspar, apatite, and opaque minerals. Nepheline and biotite are occasionally found in the groundmass.

Geochemical Characteristics of HIMU Basalts

Major element, trace'element, and modal compositions of basaltic rocks from each island are listed in Table 1, along with the CIPW-normative compositions, assuming Fe2+/(Fe 2+ + Fe 3+) = 0.85, which is the highest value observed in Polyne- sian basalts [Desonie et al., 1993]. Pb isotopic compositions are listed in Table 2. •he measured Pb isotopic ratios can be considered as initial ratios within error, because all samples are younger than 18 Ma.

Pb Isotopic Ratios

The 2ø7pb/2ønpb and 2øspb/2ø4pb ratios are plotted against 2ø6pb/2ø4pb in Figure 3, together with the field of Polynesian islands, St. Helena, and Pacific MORB. For all islands, Pb iso-

topic ratios measured in this study fall within the range of Pb isotopic ratios previously reported [Vidal eta/., 1984; Dupuy el al., 1987; Nakamura and Tatsumoto, 1988; Cheng el al., 1993; Duncan el al., 1994; Hdrnondet al., 1994; While and Duncan, 1996]. Rocks of Mangaia and Tubuai have extremely radiogenic Pb isotopic compositions (2ø6pb/2ø4pb = 21.593- 21.659 and 21.057-21.113, respectively), which are typical of HIMU value. The rocks analyzed from Rurutu, all of which belong to the old series, also have 2ø6pb/2ø4pb ratios more than 21. Rocks of Huahine, Tahiti, Ua Pou, and Hiva Oa have inter-

mediate Pb isotopic compositions. Sr and Nd isotopic data previously reported for these islands (Figure 3c) are also in- termediate compared with whole OIB variations, although they cover a wide range of compositions between MORB and EM2. Hereafter, we use the term "HIMIf' for rocks of Mangaia, Ru- rutu, and Tubuai which have much higher Pb isotopic ratios (2ø6pb/2ø4pb >20.5) and" non-HIMif' for rocks of other islands which have intermediate Pb, Sr, and Nd isotopes.

Petrography

Samples investigated in this study are basalts with phe- nocrysts less than 25 vol %. Rocks with cumulative textures, such as ankaramites, and highly under-saturated rocks, such as nephelinites, are not included The samples are all alkalic rocks as defined by Macdonald and Katsura [1964], although some samples have normative hypersthene (Table 1). •he sil- ica versus total alkalis diagram (Figure 2) shows that some al- kalic rocks plot in the tholeiitic field of the Hawaiian rocks.

In all islands investigated, alkalic basalts generally have porphyritic texture with olivine and clinopyroxene phe- nocrysts, whereas some rocks are aphyric. Rocks with rela-

Major and Trace Elements

Figure 4 shows the MgO versus major and trace element diagrams of basaltic rocks for the Polynesian islands, along with HIMUbasalts from St. Helena. Chondrite-normalized REE

abundances are shown in Figure 5. HIMU basalts have clearly higher CaO than non-HIMU basalts at any given MgO content. At higher MgO contents (>10 wt %), HIMU basalts have higher Fe203* and MnO and lower SiO 2, K20, and P205 than non-HIMUbasalts. HIMU basalts also have lower Rb, Sr, Ba, Pb, Th, Zr, Y, and Ni contents than non-HIMU basalts, whereas Nb concentrations of HIMU basalts are similar to

those of non-HIMU basalts. HIMU and non-HIMU basalts have

8088 KOGISO ET AL.' HIGH g (HIMU) • ISLAND BASALTS

Table 1. Chemical, CIPW Norm, and Modal Compositions of Polynesian Basalts

Island

Sample RRT-C6 RRT-C7 RRT-C8 ,

SiO2 wt % 44.04 43.03 44.52 TiO2 3.14 2.62 3.16 A1203 12.62 10.23 12.80 Fe203* 13.52 14.27 13.67 MnO 0.18 0.19 0.18

MgO 7.98 13.12 7.74 CaO 13.04 11.62 13.43

Na20 2.21 1.76 2.22 K20 0.40 0.30 0.44 P205 0.38 0.32 0.39 TOTAL 97.51 97.46 98.55

Rurutu

RRT-C9 RRT-C10 RRT-C12 RRT-C14 RRT-C15 RRT-C28

44.26 44.08 42.49 42.90 44.68

3.19 3.15 2.10 2.21 3.18

12.97 12.69 9.55 10.01 12.87

13.56 13.66 14.00 14.00 13.72

0.18 0.17 0.18 0.18 0.19 7.46 8.05 16.59 15.63 7.88

13.30 12.98 10.59 10.97 13.40

2.27 2.32 1.41 1.56 2.07

0.52 0.65 0.23 0.23 0.48 0.40 0.39 0.27 0.28 0.40

98.11 98.14 97.41 97.97 98.87

RRT-C43

FeO*/MgO 1.52 0.98 1.59 Na20 + K20 2.61 2.06 2.66

1.64 1.53 0.76 0.81 1.57

2.79 2.97 1.64 1.79 2.55

Rb ppm 7 5 9 Sr 429 348 496

Ba 150 122 170 Nb 38 31 38 Zr 196 161 197 Y 21 17 21 Ni 113 296 114

Pb 2.1 1.9 1.5 Th 3.0 2.0 2.4

9 11 5 4 8

469 424 316 333 425

177 176 104 118 134

38 39 23 24 38

199 197 128 135 197 21 22 15 16 21

105 117 461 411 118 1.6 1.8 2.0 0.7 1.6

2.9 3.0 1.4 1.3 2.5

La

Ce

Pr

Nd

Sm

Eu

Gd

Ho

Er

Yb

Lu

or wt % 2.4 1.8 2.6 ab 14.7 11.7 14.3 an 23.3 19.2 23.7 ne 2.1 1.7 2.4 diwo 16.2 15.2 16.9 dien 9.5 9.8 9.7 difs 6.0 4.4 6.4

hyen hyfs olfo 7.3 16.0 6.7

olfa 5.1 7.8 4.9 mt 2.6 2.8 2.7 il 6.0 5.0 6.0

ap 0.9 0.7 0.9

Olivine vol % 3.7 16.7 5.0

Clinopyroxene 3.3 7.3 4.6 Plagioclase - - - Opaque - - - Biotite - - -

Groundmass 93.0 76.0 90.4

43.12 44.52

2.21 2.72

10.53 14.47

13.61 12.71

0.17 0.17

13.89 5.64

10.87 12.81

1.39 2.96

0.39 0.62

0.25 0.34

96.43 96.96

0.88 2.03

1.78 3.58

7 12

272 435

114 146

21 27

124 170

16 21

320 69

0.4 1.0

1.4 2.4

3.1 3.8 1.4 1.4 2.8 2.3 3.6 13.9 12.6 10.1 10.5 14.6 11.8 15.1 23.7 22.3 19.1 19.6 24.4 21.4 24.4

2.9 3.8 1.0 1.5 1.5 5.3

16.6 16.5 13.2 13.8 16.5 12.9 15.4 9.4 9.6 9.0 9.2 9.5 8.5 8.0 6.4 6.1 3.3 3.5 6.2 3.5 7.0

1.8

0.8

6.4 7.3 22.7 20.8 7.1 17.0 4.2

4.8 5.1 9.0 8.7 5.1 7.7 4.1 2.7 2.7 2.7 2.7 2.7 2.7 2.5 6.1 6.0 4.0 4.2 6.0 4.2 5.2 0.9 0.9 0.6 0.7 0.9 0.6 0.8

1.7 2.0 22.2 21.3 0.7 21.3 0.0 5.8 5.2 0.8 2.2 3.9 2.6 0.5

...... 0.4

- - - 0.3 - - -

92.5 92.8 77.0 76.2 95.4 76.1 99.1

similar REE patterns, with strong enrichment in light REE relative to heavy REE (Figure 5). These features of Polynesian HIMU basalts are similar to those of St. Helena [Chaffey et al., 1989].

For both HIMU and non-HIMU basalts, samples with higher MgO contents (>10 wt %) plot roughly along olivine frac-

tionation curves on the MgOvariation diagrams (Figure 4) as follows: with decreasing MgO, CaOcontents increase; Fe203* and MnO contents and A1203/CaO ratios remain almost con- stant; concentrations of SiO2, TiO2, A1203, Na20, m20, P205, and incompatible trace elements increase; and Ni concentra- tions decrease. These correlations suggest that the chemical

KOGISO ET AL.: HIGH g (H1MU) OCEAN ISLAND BASALTS 8089

Table 1. (continued) ,

Island Rurutu Tubuai

Sample RRT-C44 RRT-C51

SiO2 wt % 43.73 42.18 TiO2 2.73 2.11 A1203 14.33 9.09 F½203' 12.73 14.16 MnO 0.17 0.18

MgO 5.77 17.29 CaO 12.73 10.42

Na20 2.99 1.12 K20 0.62 0.38 P205 0.33 0.26

TOTAL 96.13 97.19

F½O*/MgO !.99 0.74 Na20 + K20 3.61 1.50

Rb ppm 12 8 Sr 426 271 Ba 142 122

Nb 27 23

Zr 169 126

Y 20 15

Ni 68 448

Pb 1.8 1.9

Th 1.7 1.4

La 22.40 18.30

Ce 49.21 40.22

Pr 5.81 4.61

Nd 22.30 17.70

Sm 5.18 3.99

Eu 1.78 1.36

Gd 5.28 4.06

Dy 3.94 2.93 Ho 0.83 0.59

Er 1.93 1.40

Yb 1.64 1.15

Lu 0.26 0.18

or wt % 3.6 2.3

ab 13.4 8.8

an 23.9 18.7

ne 6.4 0.4

diwo 15.5 13.1

dien 8.1 8.9

difs 6.9 3.1

hyen hyfs olfo 4.4 23.9

olfa 4.1 9.3

mt 2.5 2.8

il 5.2 4.0

ap 0.8 0.6

Olivine vol % 0.2 21.2

Clinopyroxene 0.1 0.5 Plagioclase 0.3 - Opaque - - Biotite - -

Groundmass 99.4 78.3

RRT-C64 RRT-C65

45.10

2.70

13.85

13.40

0.17

8.07

12.65

2.26

0.51

0.33

99.04

1.49

2.77

10

410

145

27

168

20

119

1.5

2.3

TBA-C8 TBA-C10 TBA-Cll TBA-C54 TBA-C57 TBA-C61

3.0

15.9

26.1

1.7

14.4

8.3

5.4

44.12 42.99 43.09 41.59 43.09 43.17 43.16 2.35 2.46 2.35 3.21 3.16 3.18 2.72

10.14 11.67 9.79 15.00 14.92 15.07 11.09 13.19 13.23 13.82 14.44 14.61 14.72 12.86 0.17 0.19 0.19 0.22 0.20 0.20 0.17

14.66 11.67 14.95 5.36 5.30 5.17 9.33 11.99 12.73 11.35 11.31 11.94 11.94 14.98

1.74 2.44 1.77 5.12 3.46 3.46 2.05 0.27 0.75 0.58 1.56 1.05 1.03 0.77 0.32 0.37 0.32 0.78 0.57 0.57 0.45

98.95 98.50 98.21 98.59 98.30 98.51 97.58

0.81 1.02 0.83 0.83 2.48 2.56 1.24 2.01 3.19 2.35 2.35 4.51 4.49 2.82

6 30 15 52 25 23 24 410 473 378 925 678 669 574 146 258 173 483 379 392 398 26 48 35 107 69 69 60

156 182 157 343 235 237 205 17 21 18 30 29 39 21

327 223 343 43 49 52 115 1.3 3.0 2.1 4.9 4.1 3.0 1.9 1.3 4.5 3.6 11.9 6.6 6.8 5.1

38.20 28.00

79.87 58.77

8.95 6.57

32.60 24.60

6.51 5.29

2.14 1.74

6.11 5.08

4.27 3.64

0.87 0.75

1.96 1.72

1.63 1.44

0.26 0.23

1.6 4.5 3.4 9.2 6.2 6.1 4.5 11.1 4.3 8.0 1.3 ! 0.6 11.0 2.9 19.1 18.7 17.1 13.3 22.1 22.5 18.8 2.0 8.8 3.8 22.7 10.1 9.9 7.8

16.0 17.5 15.5 15.7 14.0 13.8 21.9 10.8 11.2 10.4 7.6 6.7 6.5 13.5 4.0 5.2 4.0 7.9 7.1 7.1 7.2

8.3 18.0 12.5 18.8 4.0 4.6 4.5 6.8 6.0 7.3 6.4 8.0 4.6 5.4 5.4 4.0 2.6 2.6 2.6 2.7 2.8 2.9 2.9 2.5 5.1 4.5 4.7 4.5 6.1 6.0 6.0 5.2 0.8 0.7 0.9 0.8 1.8 1.3 1.3 1.0

3.5 17.3 8.6 15.6 - 1.2 0.7 2.9 0.3 10.4 0.2 0.1 - 3.9 2.4 15.8

- - 0.3 0.1 ....

96.2 72.3 90.9 84.2 100.0 94.9 96.9 81.3

variations of HIMU and non-HIMU basalts at higher MgO con- tents are mainly controlled by olivine fractionation and cli- nopyroxene fractionation plays a minor role. Some samples do not plot along olivine fractionation curves at MgO >13 wt % on the MgO-Fe203' diagram (Figure 4a), which may be a result of accumulation of olivine phenocrysts and xenocrysts.

Increases of T102, A1203, and Sr with decreasing MgO suggest that fractionation of plagioclase and titanomagnetite is not significant.

These observations indicate that the differences in the

chemical composition at MgO >10 wt % between HIMU and non-HIMUbasalts may reflect differences in that of their pri-

8090 KOGISO ET AL.' HIGH g (HIMU) • ISLAND BASALTS

Table 1. (continued) ,

Island

Sample

SiO2 wt % TiO2 A1203 Fe•O3* MnO

MgO CaO

Na20 K20 P2Os TOTAL

Tubuai

TBA-C68 TBA-C70 TBA-C72

42.83 45.00 42.23 3.11 2.35 2.51

12.85 11.86 9.94 13.47 12.33 13.59 0.19 0.17 0.18 8.14 8.79 14.11

13.58 13.84 12.82 2.38 2.35 1.53 0.87 0.58 0.58 0.47 0.37 0.36

97.89 97.64 97.85

FeO*/MgO 1.49 1.26 0.87 Na20 + K20 3.25 2.93 2.11

Rb ppm 22 Sr 690 Ba 268 Nb 55 Zr 215 Y 23

Ni 97 Pb 2.2 Th 5.9

La 45.50

Ce 93.28 Pr 10.18

Nd 36.60 Sm 7.43 Eu 2.44 Gd 7.01

Dy 4.81 Ho 0.96 Er 2.19 Yb 1.79 Lu 0.28

13 15

416 421

188 247

38 43

163 161

20 19

112 332 2.0 2.6

3.3 3.2

or wt% 5.1 3.5 3.4 ab 6.1 11.4 3.4 an 21.8 20.1 18.5 ne 7.6 4.6 5.2 diwo 17.7 19.3 17.9 dien 10.4 11.6 11.9 difs 6.5 6.6 4.7

hyen hyfs olfo 6.9 7.2 16.3 olfa 4.7 4.5 7.0 mt 2.6 2.4 2.7 il 5.9 4.5 4.8

ap 1.1 0.8 0.8

Olivine vol % 2.0 3.1 19.8 Clinopyroxene 0.7 4.8 3.7 Plagioclase - - - Opaque - - 0.1 Biotite - - -

Groundmass 97.3 92.1 76.4

,

Mangaia

MGA-C9 MGA-C10 MGA-C14 MGA-C16 MGA-C17 MGA-C27 MGA-C30

42.51 43.19 42.38 42.87 44.07 2.81 2.96 3.30 3.31 3.39

11.47 12.15 14.11 14.20 14.14 13.39 13.44 14.11 14.03 14.35 0.24 0.18 0.24 0.19 0.18 9.80 8.92 6.16 5.94 5.80

13.30 13.64 12.36 12.43 12.78 2.14 2.28 2.95 3.10 3.05 0.67 0.68 0.88 0.90 0.70 0.37 0.38 0.44 0.45 0.41

96.70 97.82 96.93 97.42 98.87

1.23 1.36 2.06 2.13 2.23 2.81 2.96 3.83 4.00 3.75

14 14 18 18 15 415 418 522 524 457 208 211 244 243 196

35 37 46 46 38 180 190 208 210 206 21 21 25 25 24

160 123 51 52 54 1.8 1.4 2.6 2.7 2.1 3.1 2.8 4.3 3.7 3.0

42.09 43.39

3.70 2.92 14.21 11.39

14.94 13.18 0.19 0.18

7.30 10.56 12.18 13.39

3.01 2.36

0.98 0.77

0.39 0.40

98.99 98.54

1.84 1.12

3.99 3.13

17

514

222

43

194

24

85

2.4

3.9

33.40

73.53

8.75

33.50

7.01

2.37

6.88

4.95

1.01

2.27

1.89

0.29

16

452

208

41

190

21

161

1.4

3.0

3.9 4.0 5.2 5.3 4.1 5.8 4.6 6.2 7.1 9.4 10.0 13.1 6.2 5.5

19.7 20.9 22.7 22.2 22.8 22.4 18.2 6.5 6.6 8.4 8.8 6.9 10.4 7.8

18.3 18.5 14.9 15.3 15.8 14.8 19.0 Jl.2 11.1 7.8 7.9 8.0 8.2 12.1 6.0 6.4 6.7 7.0 7.4 6.1 5.8

9.2 7.8 5.3 4.8 4.5 7.0 10.0 5.5 5.0 5.0 4.7 4.5 5.8 5.2 2.6 2.6 2.8 2.7 2.8 2.9 2.6 5.3 5.6 6.3 6.3 6.4 7.0 5.6 0.8 0.9 1.0 1.0 0.9 0.9 0.9

10.5 3.0 - - - 5.6 6.4 3.2 8.4 0.8 0.0 - 0.7 15.4

0.1 ......

86.2 88.6 99.2 100.0 100.0 93.7 78.2

mary magmas. Olivine and clinopyroxene accumulation does not affect the differences in Fe203* and CaO contents between HIMU and non-HIMU, because HIMU and non-HiMU lavas have similar modal compositions of olivine at higher MgO (HIMU, 15-22 vol %; non-HIMU, 12-28 vol % at MgO >13 wt %; Table 1) and CaO contents do not correlate with modal

composition of clinopyroxene (Table 1). Differences in SiOn, I•O, P20•, and trace element concentrations at the same MgO content cannot be generated by olivine fractionation from single magma composition, became these elements are not fractionated relative to each during olivine crystallization [e.g., Green, 1994]. It thus appears that primary magmas of

KOGISO ET AL.: HIGH g (HIMU) OCEAN ISLAND BASALTS 8091

Table 1. (continued)

Island Mangaia

Sample MGA-C32 MGA-C34

SiO2 wt % 43.28 44.19 TiO2 2.98 3.15 A1203 11.69 13.40 Fe20•* 13.42 13.51 MnO 0.18 0.17

MgO 9.91 6.84 CaO 13.33 12.87

Na20 2.60 2.91 K20 0.83 0.76 P205 0.42 0.42

TOTAL 98.64 98.22

FeO*/MgO 1.22 1.78 Na20 + K:O 3.43 3.67

Rb ppm 16 16 Sr 471 476

Ba 224 199

Nb 43 40

Zr 196 207

Y 22 23

Ni 151 77

Pb 1.8 2.2

Th 3.3 3.3

La 34.20

Ce 75.21 Pr 8.74

Nd 33.50

Sm 7.09

Eu 2.36

Gd 6.92

Dy 4.94 Ho 1.00

Er 2.31

Yb 1.89

Lu 0.28

or wt % 4.9 4.5

ab 5.2 12.2

an 17.8 21.3

ne 9.1 6.7

diwo 19.0 16.6

dien 11.8 9.2

difs 6.1 6.8

hyen hyfs olfo 9.0 5.5

olfa 5.1 4.5

mt 2.6 2.6

il 5.7 6.0

ap 1.0 1.0

Olivine vol % 7.6 0.0

Clinopyroxene 13.0 0.9 Plagioclase - - Opaque - - Biotite - -

Groundmass 79.4 99.1

Bora Bora Huahine

BB-2 BB-3 BB-6 BB-8 HHN-1 HHN-5 HHN-6 HHN-8

47.06 45.89 47.48 46.33 47.66 46.70 46.35 46.59

3.12 2.69 3.52 2.98 2.67 2.43 2.67 2.68 12.42 11.30 14.24 11.20 11.91 10.95 11.42 11.66 12.27 12.54 12.18 12.98 12.29 12.75 12.24 12.56 0.15 0.16 0.15 0.16 0.15 0.15 0.16 0.16

10.30 13.24 5.66 12.69 11.92 14.35 13.48 13.06 10.13 10.53 11.13 9.06 9.09 8.66 9.61 9.08 2.27 2.37 2.62 2.28 2.32 2.18 2.35 2.08 1.53 0.66 1.65 1.68 1.86 1.55 1.62 1.85 0.39 0.34 0.47 0.44 0.43 0.38 0.44 0.43

99.64 99.72 99.10 99.80 100.30 100.10 100.34 100.15

1.07 0.85 1.94 0.92 0.93 0.80 0.82 0.87 3.80 3.03 4.27 3.96 4.18 3.73 3.97 3.93

32 23 28 33 38 35 32 41

478 386 633 494 495 430 535 544 254 176 344 296 419 377 417 454

26 20 29 28 31 27 35 34 253 192 287 273 267 210 246 251

22 19 26 21 23 20 21 21

324 438 72 368 324 406 332 353 2.1 1.7 2.5 3.5 3.9 3.2 3.4 3.8 3.3 2.7 3.6 2.7 4.2 3.8 5.4 5.8

20.90 30.20 28.60 34.30 35.70 50.88 69.45 62.52 73.17 76.94

6.06 8.42 7.13 8.33 8.67 23.90 32.80 27.30 31.20 32.60

5.45 7.09 5.97 6.70 6.96 1.81 2.28 1.92 2.13 2.21

5.58 6.69 5.77 6.28 6.55 4.17 4.39 3.92 4.19 4.32

0.85 0.85 0.77 0.8! 0.83 2.06 1.93 1.77 1.82 1.85 1.81 1.54 1.43 1.44 1.47 0.28 0.23 0.21 0.21 0.21

9.0 3.9 9.8 9.9 11.0 9.1 9.6 10.9 19.2 17.4 22.2 17.4 19.6 17.9 14.5 16.4 19.2 18.2 22.2 15.3 16.6 15.5 15.8 17.0

1.4 1.0 0.3 2.9 0.7 11.9 13.3 12.5 11.2 10.7 10.4 12.1 10.5 7.8 9.0 7.0 7.5 7.1 7.1 8.3 7.1 3.3 3.3 5.0 2.9 2.8 2.5 2.9 2.6 0.7 0.2 0.1

0.3 0.2 0.0

12.0 16.8 4.8 16.9 15.8 20.1 17.7 17.8 5.7 6.8 3.9 7.2 6.9 7.9 6.8 7.3 2.4 2.5 2.4 2.5 2.4 2.5 2.4 2.5 5.9 5.1 6.7 5.7 5.1 4.6 5.1 5.1 0.9 0.8 1.1 1.0 i.0 0.9 1.0 1.0

14.5 28.9 0.8 20.7 16.1 19.8 17.0 18.6 3.7 - 0.8 0.1 0.7 0.1 3.3 1.5

- - - 0.3 - - 0.1 -

81.8 71.1 98.4 78.9 83.2 80.1 79.6 79.9

HIMU basalts are enriched in Fe203*, MnO, and CaO and de- pleted in SiO2, Ni, and incompatible elements, except for Nb, than those of non-HIMUbasalts. In the following section, we shall consider possible explanations for the differences in primary magma compositions and discuss the possible origin of the HIMU component.

Origin of the HIMU Component

Constraints From Major Element Characteristics

We here examine whether HIMU and non-HiMU primary magmas are generated from different sources or whether both magma types can be generated from single-source composition

8092 KOGISO ET AL.' HIGH !.t (HIMU) OCEAN ISLAND BASALTS

Table 1. (continued)

Island Tahiti Ua Pou

Sample THT-A THT-I5 THT-I6 UAP-01 UAP-02 UAP-101 UAP-117 UAP-150 UAP-167 UAP-172 ,

SiO2 wt % 42.39 43.93 44.40 45.05 42.20 42.44 43.40 43.89 43.38 43.55 TiO2 4.69 2.89 3.68 3.73 4.15 4.13 4.05 3.91 4.11 4.13 A1203 14.42 10.18 16.94 15.88 13.53 13.57 13.99 15.02 14.47 14.48 Fe•O3* 14.30 13.61 12.99 11.94 13.83 13.83 13.30 13.08 13.53 13.65 MnO 0.18 0.18 0.20 0.19 0.19 0.19 0.16 0.18 0.20 0.17 MgO 6.92 15.25 4.74 5.53 8.12 8.03 7.80 5.34 7.51 7.08 CaO 12.74 10.58 10.16 9.68 10.59 10.58 10.28 10.08 10.76 10.54 Na•O 1.64 1.70 4.12 3.36 3.50 3.59 2.75 3.99 2.39 3.65 K•O 0.63 0.59 1.20 1.03 0.67 0.64 0.94 1.00 1.01 0.57 P205 0.54 0.43 0.69 0.83 0.79 0.78 0.68 0.95 0.70 0.74 TOTAL 98.45 99.34 99.12 97.22 97.57 97.78 97.35 97.44 98.06 98.56

FeO*/MgO 1.86 0.80 2.47 1.94 1.53 1.55 1.53 2.20 1.62 1.74 Na20 + K•O 2.27 2.29 5.32 4.39 4.17 4.23 3.69 4.99 3.40 4.22

Rb ppm 6 6 110 42 80 72 80 101 50 61 Sr 816 475 925 1017 917 921 844 1091 879 885 Ba 544 322 553 706 623 621 671 695 611 581 Nb 41 30 59 64 59 59 59 74 58 58 Zr 361 246 317 317 323 325 320 381 313 317 Y 29 22 29 29 28 28 27 30 39 27 Ni 74 367 17 37 100 96 125 35 89 89 Pb 3.0 3.0 3.2 4.9 4.4 4.3 3.7 5.1 3.5 4.2 Th 4.0 3.4 4.9 8.7 7.5 8.3 7.8 10.4 8.1 7.4

La 29.50

Ce 68.92 Pr 8.24 Nd 32.20

Sm 7.08 Eu 2.28

Gd 6.80

Dy 4.53 Ho 0.87 Er 1.95 Yb 1.53 Lu 0.22

or wt % 3.7 3.5 7.1 6.1 4.0 3.8 5.5 5.9 6.0 3.4 ab 13.9 13.0 18.1 25.7 14.8 15.4 18.8 20.8 17.9 19.2 an 30.1 18.4 24.2 25.2 19.2 19.0 23.0 20.1 25.8 21.4 ne 0.7 9.1 i.5 8.0 8.1 2.4 7.0 1.2 6.3 diwo 12.3 13.1 9.0 7.3 11.8 11.8 9.8 9.9 9.6 10.9 dien 7.2 8.9 4.6 4.1 7.1 7.1 6.0 5.3 5.7 6.3 difs 4.6 3.1 4.3 2.9 4.0 4.1 3.3 4.3 3.4 4.0

hyen 0.5 hyfs 0.3 olfo 6.7 20.4 5.1 6.8 9.2 9.0 9.4 5.6 9.1 7.9 olfa 4.8 7.8 5.3 5.2 5.7 5.7 5.8 5.0 5.9 5.5 mt 2.8 2.7 2.5 2.3 2.7 2.7 2.6 2.6 2.6 2.7 il 8.9 5.5 7.0 7.1 7.9 7.8 7.7 7.4 7.8 7.8

ap 1.2 1.0 1.6 1.9 1.8 1.8 1.6 2.2 1.6 1.7 Olivine vol % 2.2 21.0 - 1.5 4.2 4.0 9.8 1.5 3.5 5.0 Clinopyroxene 4.2 - - 8.5 4.7 7.7 8.9 10.8 6.0 7.5 Plagioclase .......... Opaque - 0.8 - 1.0 0.7 0.7 0.6 2.2 1.3 0.5 Biotite .........

Groundmass 93.6 78.2 100.0 89.0 90.4 87.6 80.7 85.5 89.2 87.0

under different physical conditions, on the basis of the results of melting experiments on mantle peridotite. Recent studies of peridotite melting have revealed the detailed relationship be- tween melt compositions and phyõical conditions of melting [e.g., Tak ah as h i and Kus h i ro , 1983; Fal l o o n eta/., 1988; Kinzler and Grove, 1992a, b; Hirose and Kushiro, 1993;

Kushiro, 1996]. Although these studies were restricted to batch melting under dry conditions, some important features of chemical variations during partial melting were obtained. (1) SiO 2 contents of melt decrease with increasing pressure but de- pend little on the degree of melting. (2) MgO and FeO* con- tents increase with increasing pressure. (3) TiO2, Na20, and

KOGISO ET AL.: HIGH I.t (HIMU) OCEAN ISLAND BASALTS 8093

Table 1. (continued)

Island Ua Pou Hiva Oa

Sample UAP- 185 UAP-211 UAP-301 UAP-302 UAP-308 UAP-309 UAP-361 UAP-370 HVO-BB HVO-BD

SiO2 wt % 43.16 44.19 47.86 45.67 43.55 44.21 43.24 44.45 46.87 47.84 TiO2 3.28 2.82 2.85 3.74 2.90 2.78 3.92 4.10 3.51 3.90 A1203 14.35 13.29 18.61 16.23 12.01 12.59 14.30 14.85 12.47 14.44 Fe203* 13.73 12.75 8.66 12.00 13.16 13.12 13.24 12.98 12.64 11.99 MnO 0.18 0.18 0.18 0.19 0.18 0.18 0.19 0.17 0.16 0.15

MgO 8.30 10.55 3.70 5.55 12.12 11.93 7.76 6.05 10.14 6.01 CaO 10.76 10.52 7.55 9.69 12.30 10.36 10.47 10.56 9.75 10.47

Na20 2.78 3.05 5.40 3.81 2.93 3.43 2.63 3.86 2.54 2.81 K20 1.64 1.37 2.19 1.15 0.58 0.64 1.84 0.86 1.52 1.70 P205 0.56 0.49 0.89 0.82 0.41 0.48 0.88 0.88 0.46 0.52

TOTAL 98.74 99.21 97.89 98.85 100.14 99.72 98.47 98.76 100.06 99.83

FeO*/MgO 1.49 1.09 2.11 1.95 0.98 0.99 1.54 1.93 1.12 1.80 Na20 + K20 4.42 4.42 7.59 4.96 3.51 4.07 4.47 4.72 4.06 4.51

Rb ppm 73 42 83 44 11 14 101 47 29 39 Sr 834 728 1180 1045 642 717 1051 986 565 618

Ba 564 525 807 680 400 506 631 672 281 294

Nb 50 47 84 64 37 44 62 67 29 32

Zr 221 228 402 312 178 209 321 360 277 305 Y 23 21 31 32 19 20 29 29 25 30

Ni 81 217 3 38 227 267 91 73 238 64

Pb 3.3 4.4 7.3 4.2 2.8 4.1 4.2 4.9 2.7 2.0

Th 6.2 5.9 11.7 8.8 3.5 6.0 8.5 10.1 3.0 3.7

La 47.50 30.90 31.60 32.90

Ce 95.63 66.13 72.90 76.76

Pr 10.27 7.58 8.75 9.32

Nd 36.10 28.90 34.40 37.20

Sm 7.02 6.18 7.73 8.73

Eu 2.31 2.10 2.50 2.86

Gd 6.55 5.96 7.60 8.74

Dy 4.51 3.94 5.26 6.14 Ho 0.90 0.76 1.04 1.22

Er 2.03 1.69 2.35 2.71

Yb 1.65 1.29 1.84 2.11

Lu 0.25 0.19 0.27 0.31

or wt % 9.7 8.1 13.0 6.8 3.4 3.8 10.9 5.1 9.0 10.0

ab 8.7 9.9 26.8 24.3 6.4 12.8 12.0 20.7 20.7 23.8

an 21.9 18.5 20.1 23.8 17.9 17.1 21.8 20.7 18.1 21.8

ne 8.0 8.6 10.2 4.3 9.9 8.8 5.6 6.5 0.4

diwo 11.6 12.7 4.8 7.9 16.9 13.0 10.2 10.9 11.3 11.2

dien 6.9 8.1 2.7 4.5 11.1 8.5 6.2 6.2 7.4 6.6

difs 4.2 3.8 2.0 3.1 4.6 3.6 3.5 4.2 3.2 4.1

hyen 0.7 hyfs 0.4 olfo 9.7 12.7 4.6 6.6 13.4 14.9 9.2 6.2 12.5 5.4

olfa 6.6 6.5 3.7 5.1 6.2 7.0 5.8 4.7 5.9 3.7

mt 2.7 2.5 1.7 2.3 2.6 2.6 2.6 2.5 2.5 2.3

il 6.2 5.4 5.4 7.1 5.5 5.3 7.4 7.8 6.7 7.4

ap 1.3 1.1 2.1 1.9 1.0 1.1 2.0 2.0 1.1 1.2

Olivine vol % 8.3 13.2 - 2.1 11.5 13.4 3.5 3.4 15.3 0.6

Clinopyroxene 5.6 0.8 - 5.7 0.2 2.2 5.6 7.5 0.3 2.4 Plagioclase .......... Opaque 0.7 0.1 - 1.4 0.0 0.1 0.9 1.6 0.1 - Biotite ..... 0.3 ....

Groundmass 85.4 85.9 100.0 90.8 88.3 84.0 90.0 87.5 84.3 97.0 , ,

K20 behave as incompatible elements, their contents decreas- ing with increasing degree of partial melting. (4)CaO contents increase with increasing degree of partial melting until cli- nopyroxene disappears and then decrease. (5) A1203 contents decrease with increasing degree of partial melting, even if gar- net exist in the residue.

It appears that the chemical differences between HIMU magmas and non-HIMU magmas in Polynesia cannot be pro- duced from the same source composition. In order to produce HIMU magmas with lower K20 and incompatible trace ele- ments than those of non-HIMU from the same source compo- sition, primary magmas of HIMU would have to be generated at

8094 KOGISO ET AL.' HIGH [.t (HIMU) OCEAN ISLAND BASALTS

Table 1. (continued)

Island Hiva Oa

Sample HVO-BH HVO-BK HVO-O HVO-Q HVO-R6 HVO-X

Si02 wt % 47.85 45.33 48.39 47.48 49.20 44.92 TiO2 4.09 3.93 4.27 3.13 3.79 3.67 A1203 14.11 12.49 14.53 12.27 15.81 12.37 F½203' 12.37 13.76 10.71 12.58 12.04 13.40 MnO 0.14 0.15 0.11 0.15 0.16 0.18

MgO 6.12 11.03 5.35 11.19 4.23 10.98 CaO 9.88 9.01 11.04 8.42 8.44 9.13

Na•O 3.15 2.51 3.27 2.92 3.57 2.36 K•O 1.37 1.08 1.12 1.16 2.38 1.64 P•O5 0.57 0.53 0.53 0.52 0.64 0.50

TOTAL 99.65 99.82 99.32 99.82 100.26 99.15

FeO*/MgO 1.82 1.12 1.80 1.01 2.56 1.10 Na:O + K20 4.52 3.59 4.39 4.08 5.95 4.00

Rb ppm 27 22 12 21 61 40 Sr 679 638 652 574 687 607

Ba 242 200 187 210 384 293

Nb 41 39 39 35 41 34

Zr 345 323 345 328 374 296

Y 27 26 28 25 34 26

Ni 98 295 83 320 25 254

Pb 2.7 2.2 1.9 1.9 4.0 3.1

Th 3.9 3.1 3.0 4.6 5.7 4.1

La 36.50 38.00

Ce 83.63 85.62

Pr 10.27 10.00

Nd 41.00 38.80

Sm 9.31 8.54

Eu 3.15 2.84

Gd 9.17 8.17

Dy 6.19 5.37 Ho 1.21 1.03

Er 2.70 2.30

Yb 2.13 1.84

Lu 0.31 0.27

or wt % 8.1 6.4 6.6 6.8 14.0 9.7

ab 26.6 21.3 27.7 24.7 29.4 16.7

an 20.3 19.6 21.6 16.9 20.1 18.3

ne 0.4 1.8

diwo 10.4 9.0 12.4 9.0 7.4 9.9

dien 6.1 5.9 7.6 5.9 3.7 6.5

difs 3.8 2.5 4.0 2.4 3.4 2.8

hyen 1.6 0.7 1.3 2.2 hyfs 1.0 0.3 0.7 0.9 olfo 5.3 14.6 3.1 13.9 4.8 14.6

olfa 3.6 6.8 1.8 6.2 4.8 6.8

mt 2.4 2.7 2.1 2.5 2.4 2.6

il 7.8 7.5 8.1 5.9 7.2 7.0

ap 1.3 1.2 1.2 1.2 1.5 1.2

Olivine vol % 1.0 15.7 0.1 16.8 - 16.7

Clinopyroxene 2.5 3.2 1.7 3.5 - 3.0 Plagioclase - 0.3 - 0.3 - 1.2 Opaque 0.1 0.4 - 0.2 - - Biotite ......

Groundmass 96.4 80.4 98.2 79.2 100.0 79.1

F½203' and FeO* denote total iron as F½203 and FeO, respectively. Dash denote not detected.

larger degrees of partial melting than those of non-HIMU. However, larger degrees of partial melting lead to higher Ni concentrations in melts, because Ni is highly compatible in mantle phases, which is contrary to the observation that pri-

mary magmas of HIMU have lower Ni concentrations than non-HIMU (Figure 4).

Alternatively, the differences between HIMU and non-HIMU primary magma compositions may be explained by differences

KOGISO ET AL.' HIGH g (HIMU) OCEAN ISLAND BASALTS 8095

z•' 5

I1

ß Mangala ß Rurutu ß Tubuai a Huahine o Bora Bora v Tahiti [] Ua Pou n Hiva Oa

49 50

Figure 2. SiO2 versus total alkalis diagram for Polynesian basalts. The solid line is the boundary between tholeiitic and alkalic fields for Hawaiian rocks from Macdonald and Katsura

[19641.

in source composition: HIMU magmas may be generated from a more fertile source than non-HIMU magmas. Melting ex- periments on compositionally different peridotites [Hirose and Kushiro, 1993] have suggested that partial melts of a more fertile (SiO 2- and FeO-enriched and MgO-depleted) peridotite are more enriched in FeO but depleted in SiO2 than those of a more refractory (SiO 2- and FeO-depleted and MgO-enriched) peridotite at the same pressure (Figure 6), whereas MgO con- tents are not so affected by source compositions but depend on temperature. Therefore the higher Fe203* and lower SiO 2 of HIMU primary magmas than non-HIMU magmas at the same MgO suggest that the mantle source of HIMU could be more fertile than those of non-HIMU. The differences in SiO2 con- tents between HIMU and non-HIMU are not so systematic as CaO, consistent with the experiments in which partial melts of a fertile peridotite have slightly lower SiO2 contents than those of the refractory peridotite (Figure 6). High CaO con- tents of HIMU magmas than non-HIMU magmas also suggest that the HIMU source is enriched in CaO compared with non- HIMU sources, because partial melts from more CaO-rich peri- dotite are expected to have higher CaO contents (Figure 6). Concentrations of Ni in HIMU magmas are also lower than in non-HIMU magmas, indicating that the HIMU source is poorer in Ni than the non-HIMU sources. It may be suggested that the HIMU source involves a basaltic component, because basalts generally have higher SiO2, FeO, and CaO and lower Ni than mantle peridotite. Figure 6 demonstrates that HIMU basalts have lower SiO2 and higher Fe203* contents than partial melts of a fertile peridotite (HK-66: Mg # [Mg/(Mg + Fe2+)] = 85), implying that the HIMU source may be much more fertile than HK-66.

On the other hand, A1203 contents in magmas do not di- rectly reflect the source signature. A120 • contents of partial melts generated at higher pressures (>2.5 GPa) are not sensi- tive to the source composition [Hirose and Kushiro, 1993; Kushiro, 1996]. Strongly light REE-enrichedpattems for both HIMU and non-HIMU basalts (Figure 5) indicate that HIMU and non-HIMU magmas are generated in the presence of garnet in the residue. It follows that similar A120 • contents in HIMU and non-HIMU magmas do not require that the HIMU source has similar A120 • contents compared with the non-HIMU sources. It is possible that partial melts of fertile peridotite, enriched in

A1203, would have A1203 contents similar to partial melts of refractory peridotite.

Chauvel et al. [1992] attributed the higher Fe contents of HIMU basalts to higher pressure and temperature of partial melting than other OIB magmas on the basis of the model of Albarbde [1992], which does not take into account the fertility of the source. However, the experiments by Hirose and Kushiro [1993] demonstrated that the composition of partial melts is significantly affected by the source compositions. The above discussions demonstrate that the differences in major element composition between HIMU and non-HiMU magmas are ex- plained by differences in the fertility of the source. It is thus plausible that HIMU magmas are generated from a more fertile source, i.e., a source more enriched in a basaltic component than non-HIMU magmas.

Recycling of Subducted Oceanic Crust

The above discussion points to a conclusion that a basaltic component is incorporated in the source of HIMU magmas. A possible mechanism causing such incorporation is recycling of ancient subducted oceanic crest [e.g., Hofmann and White, 1982; Dupuy et al., 1988, 1989; Weaver, 1991; Chauvel et al., 1992; Hauri and Hart, 1993 ]. Recycling of the subducted oce- anic crust, with basaltic composition, can explain not only the major element characteristics of HIMU suites, but also their isotopic characteristics [Brenan et al., 1995; Kogiso et a/., 1997]. The extremely high Pb isotopic ratios of HIMU can be produced by the involvement of subducted oceanic crust, be- cause U/Pb and Th/Pb ratios of the subducted oceanic crust in-

crease during dehydration processes in subduction zones. Sr isotopic ratios of the HIMU component can also be produced by mixing of subducted oceanic crust with less depleted mantle [Brenan et al., 1995].

Constraints From Incompatible Elements

Lower concentrations of K20, P205, and incompatible trace elements, except for Nb, in HIMU magmas than in non-HIMU magmas suggest that the HIMU source is less enriched in in- compatible, elements relative to non-HIMU sources. Non- HIMUbasalts have intermediate Sr, Nd, andPb isotopic ratios,

Table 2. Pb Isotopic Compositions for Polynesian Basalts

2ø•pb?ø4pb 2ø7pb?ø4pb 2øspb?ø4pb

RRT-C10 21.082 +0.004 15.764 _+0.004 40.286 +0.008

RRT-C65 21.068 +0.003 15.756 +0.003 40.292 _+0.006

TBA-C8 21.057 +0.019 15.753 +0.016 40.294 +0.038

TBA-C11 21.113 +0.005 15.738 +0.004 40.362 +0.009

MGA-C15 21.607 +0.004 15.796 +0.004 40.484 +0.009

MGA-C27 21.659 +0.006 15.832 +0.006 40.630 _+0.013

MGA-C34 21.593 +0.019 15.825 _+0.009 40.572 +0.036

HHN-8 19.096 +0.003 15.611 +0.002 38.941 +0.006

THT-I5 19.087 +0.003 15.583 +0.002 38.788 +0.006

UAP-301 19.266 +0.002 15.578 +0.002 39.173 +0.006

HVO-BD 18.901 +0.004 15.556 +0.004 38.697 +0.009

HVO-R6 18.982 +0.004 15.583 +0.004 38.836 +0.011

RRT is Rurutu, TBA is Tubuai, MGA is Mangaia, HHN is Huahine, THT is Tahiti, UAP is Ua Pou, and HVO is Hiva Oa.

Correction factors for mass fractionation are 0.11% for 2ø•pb/2øapb and 2ø7pb?øapb and 0.13% for 2øSpb/2øapb.

Values for NIST-981 standards are: 2ø•pb?øapb = 16.937 + 0.005, 2ø7pb?øapb = 15.491 +0.008, and 2øSpb?øapb = 36.720+0.023 [Catanzaro et al. , 1968].

8096 KOGISO ET AL.' HIGH g (HIMU) OCEAN ISLAND BASALTS

15.9

15.8

15.7

15.6

15.5

41.0

40.5

40.0

39.5

39.0

38.5

38.0

(a) I I I I

EM2

Ua Pou

Bora Bor• hine

/Pacific MORB

St. Helena

I I

•i HIMU

ß Mangaia

© Rurutu [] Tubuai a Huahine v Tahiti [] Ua Pou [] Hiva Oa

Tubuaf•,•

I i i i

- Ua Pou

Bora

/• 'Huahine • ahiti

'acific MORB

18.5 19.0 19.5 20.0 20.5 21.0 21.5 22.0

206pb/204pb 0.5134

0.5132

0.5130

0.5128

0.5126

0.5124

(c)

va Oa

- Mangai• St. HelenaN N

EM1 I I I I I 0.5122

0.702 0.703 0.704 0.705 0.706 0.707 0.708

87Sr/86Sr

Figure 3. Isotopic ratios of Polynesian basalts, showing (a) 2ø6pb/2ø4pb-2ø7pb/2ø4pb, and (b) 2ø6pb/2ø4pb- 2øspb/2ø4pb diagrams. Symbols, as shown in inset, are values measured in this study, and data fields are those of literature. (c)Sr and Ndisotopic ratios previously reported for the studiedislands and St. Helena. Data are from Vidal et al. [1984], Dupuy et al. [1987], Nakamuraand Tatsumoto [1988], Chaffey et al. [1989], Cheng et al. [1993], Hauri and Hart [1993], White [1993], Duncan et al. [1994], Hdmond et al. [1994], White and Duncan [1996] and H. G. Barsczus unpublished data, 1995.

KOGISO ET AL.' HIGH !.t (HIMU) OCEAN ISLAND BASALTS 8097

50

48

• 46

5.0

4.5

4.0

:3.5

:3.0 2.5

18

16

lO

16

15

12

ß Mangaia o BoraBora ß Rurutu v Tahiti ß Tubuai [] UaPou a Huahine [3 HivaOa

0.30

0.25 ß ß

0.20 • • '

0.15

0.10 , i i i , •

4 6 8 10 12 14 16 18

MgO (wt. %)

16 15

- • 14 ø• 13

I i I I I ! ß (a)

•,• o A 9 2aoa

4.5 I I I I I :

4.0 •• 3.5 mm N • 3.0 $ •o m 2.5 - - 2.0 ß m•*ø a a•*a _

1.5 • • • 2.5 I• • I I I I I 2.0

•0 ø • o 0 •

1.0 ß • - m • l

0.5 {

1.0 , • ---

0.8 •

0.6 • • v • o• mo• _ _ •

2.0 .... • • •{ 1.8

1.6

1.4

1.2 1.0 0.8

0.6

v •

4 6 8 10 12 14 16

MgO (wt. %) 18

Figure 4. MgO variation diagrams of (a) major elements and (b) trace elements for Polynesian basalts, compared with HIMU basalts. from St. Helena [Chaffey et al., 1989]. Symbols are shown in inset. At high MgO contents (>10 wt %), HIMU basalts (solid symbols) are characterized by higher contents in Fe203*, MnO, and CaO and lower contents in SiO2, K20, P205, Rb, Sr, Ba, Pb, Th, Zr, Y, and Ni than non-HIMU basalts (open symbols). HIMU basalts from St. Helena have similar characteristics to Polynesian HIMU basalts.

8098 KOGIS6 ET AL.' HIGH It (H1MU) OCEAN ISLAND BASALTS

120

lOO

,•- 80

20

1200

1000

800

600 400

800

7OO

6OO

i i I i I i

- m mm A A A

• 500 [] '•' 400

I:• 300 [] o

[ 6 • i i i i i i

• 4

12

m o •

ß v

80

70 ß

6o 5O

40 m• •/ :30

400

350

300

250

200

150

35

30

25

20

5OO

4OO

300

2O0

100

i i i i i

Co)

i i i l i i i i

t I I I I I I 4 6 8 10 12 14 16 18

MgO (wt. %)

Figure 4. (continued)

KOGISO ET AL.' HIGH g (HIMU) OCEAN ISLAND BASALTS 8099

2OO

IO0

10

200

lOO

lO

(a) HIMU i [ i i ! i ! i i ! i i ! !

lena

(b) non-HIMU

i i i i i i i i i i i i i I i

La Ce Pr Nd Sm EuGd DyHo Er Yb Lu

Figure 5. Chondrite-normalized rare earth element (REE) abundances in (a) HIMU and (b) non-HIMU basalts, compared with HIMU basalts from St. Helena [Chaffey et al., 1989]. Symbols are the same as in Figure 4. Values of chondrite are taken from Nakarnura [ 1974].

trending towardEM2 (Figure 3), suggesting an involvement of this component. EM2 basalts are characterized by very ra- diogenic Sr isotop!c ratios and enriched incompatible ele- ments, suggesting that the EM2 source involves subducted sediments [White and Hofmann, 1982]. It seems'likely that non-HIMU sources are enriched in incompatible elements by the involvement of sedimentary components.

Some authors have suggested that EM1 and EM2 sourc_es also contain subducted oceanic crest in addition to sedimentary components [e.g., Weaver, 1991' Chauvel et al., 1992' Roy- Barman andAllbgre, 1995]. They demonstrated that addition of small amounts (-3 %)of sedimentary components to the HIMU source account for the incompatible trace element and isotope characteristics of EM2 basalts. However, the major element characteristics of HIMU and non-HIMU magmas do not support a significant involvement of basaltic components in non- HIMU sources. On the other hand, an involvement of small

amounts of sediments in the source region may not signifi- cantly affect major element compositions, such as Mg# and CaO contents of magmas, because oceanic s'ediments have very

lowFeO, MgO, and CaOcontents [McLennan, 1995]. For ex- ample, addition of 3% sediments leads to only 0.15 wt % in- crease in FeO contents. It is thus likely that non-HIMU sources contain sedimentary components without subducted oceanic Crust.

Relative Enrichment of Nb

A serious problem, however, arises when we explain rela- tive enrichment of Nb in HIMU magmas by simple addition of subducted components to the source region. HIMU basalts have systematically higher Nb/Zr and Nb/Rb ratios than non-

49

48

47

.• 46

eq 45

43

15

14

ß 13 ß -

• 12 - ß .

• 11 - z'--Q------- ß

?

8 I I I I I ! I

10 11 12 13 14 15 16 17 18

MgO (wt. %) Figure 6. Comparison of •MU and non-•MU pfim• basalts •th experimental melts of feaile pefidotite (•-66) andrefracto• pefidotite (•-1)produced at more than 2 GPa [Hirose •dK•hiro, 1993]. Symbols •e the same as in Fig•e 4. Nmbers in the experimental melts gea in,cate press•es of melting. •-66 has higher SiO2 •dFe2Oa* contents and lower CaOcontents than •-1. SiO2 and Fe2Oa* contents of •MU pfim• bas•ts are close to pa•ial melts of•-66, •d those CaO contents ge close to p•tial melts of •B-1.

8100 KOGISO El' AL.: HIGH g (HIMU) • ISLAND BASALTS

o

o.oo

subducted MORB

g '•:•!!:•O HIMU in Cook- :•/ I Society & \/": '•'•½• Austral •/ PM + 30% • ' ]:'":•':•"••:':•':'--'•:-•••• ..... .•. ;....• - .•[ subducted ..... -a ..... ['""--'"' '"'""•<':'•::•••':••--•St. H lee na/•

i

PM + 5% SD -• ..... --= ,SD• , , I

0.05 0.10 0.15 0.20 0[25 0.30

Nb/Zr

Fig ure 7. Zr/Nb-Rb/Nb diagram for HIMU and non-HiMUbasalts compared with other Polynesian basalts and St. Helena basalts. Symbols are the same as in Figure 4. PM is primitive mantle; SD is sediments. HIMU basalts are characterized by higher Nb/Zr andNb/Rb ratios than non-HIMUbasalts. The solidarrow indicates compositional change of mid-oceanic ridge basalts (MORB)during dehy .dration reaction in subduction zones [Brenan et al., 1995; Keppler, 1996; Kogiso et al., 1997]. The thick and thin stippled lines indicate compo- sitions of partial melts of a fertile mantle (70% primitive mantle + 30% dehydrated MORB: the mineral as- semblage is 30% olivine, 20% orthopyroxene, 26% clinopyroxene and 24% garnet) and partial melts of a less fertile mantle (95% primitive mantle + 5% sediments: 44% olivine, 27% orthopyroxene, 13% clinopyroxene and 16% garnet), respectively. Numbers on the lines indicate percent of partial melting. Partition coefficients used in the calculation are from Hart and Dunn [1993] and Green [1994]. Values for the primitive mantle, MORB, and sediments are from Hofmann [1988] and McLennan [ 1995]. Data sources for OIBs are the same as in Figure 3.

HiMUbasalts (Figure 7). Oceanic sediments are enriched in Rb, Nb, and Zr but have higher Nb/Zr ratios than MORB and the primitive mantle [Hofmann, 1988' McLennan, 1995]. During dehydration proces.ses in subduction zones, Rb is readily ex- tracted from subducted lithosphere, whereas Nb and Zr are not fractionatedfrom each other[Brenan et al., 1994, 1995' Kep- pler, 1996]. Thus, if the HIMU source contains subducted oce- anic crest and non-HIMU sources contain sedimentary compo- nents, then the HIMU source is expected to have lower Nb/Zr ratios than non-HIMU sources, which is inconsistent with the observation (Figure 7). It is thus difficult to explain the en- richment of Nb relative to other incompatible elements in HIMU by simple addition of the crustal component.

Differences in Nb/Zr ratios between HIMU and non-HIMU

magmas could be produced by different degrees of partial melt- ing, because Nb is much more incompatible to upper mantle phases than Zr [Green, 1994]. Garnet may fractionate Nb and Zr most efficiently among the upper mantle phases (D• = ~0.01, Dz• =-0.3). The major element characteristics of HIMU lavas, as discussed above, suggest that the HIMU source is more fertile than non-HIMU sources, implying that the HIMU source may contain more amounts of garnet than non-HIMU sources. Here we calculated Nb/Zr and Nb/Rb ratios of magmas produced from fertile and less fertile peridotites, assuming simple modal batch melting (Figure 7). A fertile peridotite (primitive mantle + 30% subducted MORB), containing 24% garnet, produces melts with high Nb/Zr ratios (-0.2) by 2% partial melting. In contrast, a less fertile peridotite (primitive

mantle + 5% sediments), containing 16% garnet, produces melts with Nb/Zr ratios as low as 0.15 by about 5% partial melting. Most non-HIMU basalts have relatively low Nb/Zr ratios (<0.15), requiring much higher degree of melting'than HIMU magmas (Figure 7). However, it seems unlikely that the more fertile HIMU source produces magmas at lower degree of melting than less fertile non-HIMU sources. If HIMU magmas were produced by much lower degree of melting than non- HIMU magmas, HIMU basalts are expected to be much more alkalic and more'silica undersaturated than non-HIMU basaltsl which is inconsistent with the observation (Figure 2). Moreo- ver, similar REE patterns in HIMU and non-HiMU basalts (Figure 5) do not support different degree of partial melting in the presence of garnet between HIMU and non-HIMU. It is thus difficult to explain the difference in Nb/Zr ratios between HIMU and non-HIMU magmas by differences in degree of par- tial melting.

Thus the relative enrichment of Nb in HIMU magmas cannot be explained by simple addition of the crustal component to •he source region nor partial melting processes in the upper mantle. This implies that other processes in the deep mantle would contribute to fractionate Nb from other incompatible elements in the HIMU source. Mantle phases which can frac- tionate Nb/Zr ratios effectively are silicate perovskites. Kato et al. [1988a, b] showed that Zr is strongly partitioned into Mg-and Ca-perovskites (partition coefficients >5), whereas partition coefficients of Nb for these phases are low (~1). It means that partial melting processes in the presence of silicate

KOGISO ET AL.: HIGH g (HIMU) • ISLAND BASALTS 8101

perovskites could result in higher Nb/Zr ratios. Recent seismic studies [Kendall and Silver, 1996; Williams and Garnero, 1996] suggested a possibility of partial melting of subducted oceanic crust in the D" layer. If it is the case, partial melts of subducted crust would "refertilize" the surrounding mantle and then produce a fertile reservoir with high Nb/Zr ratios in the lowermost mantle.

Such refertilization processes are inconsistent with other trace element characteristics of HIMU basalts. The refertilized

reservoir should be depleted in Pb and Th, which are more compatible to silicate perovskites than Zr [Kato eta/., 1988a, b], but Pb and Th are not so depleted as Zr in HIMU basalts compared with non-HIMU basalts (Figure 4). Therefore entire trace element characteristics of HIMU basalts seem to be diffi-

cult to explain by refertilization processes involving silicate perovskites at the base of the mantle. However, silicate perovskites in a basaltic system contain significant amounts of A1 [Kesson et al., 1994], which may affect partitioning of trace elements. Subducted oceanic crust in the lower mantle

may also contain other silicate phases, such as a sodic alumi- nous phase with Ca-ferrite structure [Kesson et al., 1994]. Element partitioning data for these phases are not yet con- strained. Moreover, Saxena eta/. [1996] experimentally dem- onstrated that Mg-perovskite may not be stable in the lower mantle. Further work is needed to evaluate whether trace ele-

ment characteristics of HIMU would be explained by partial melting processes in the bottom of the mantle or not.

Implications for the Dynamic Processes in the Earth's Mantle

The above discussions suggest that in addition to the in- volvement of subducted oceanic crust, some fractionation

processes in the deep mantle are required to explain the trace element characteristics of HIMU. This implies that the sub- ducted oceanic crust may convect at whole-mantle scale. This is supported by several additional observations. Density rela- tionships between the subducted lithosphere and the mantle [Ringwood and Irifune, 1988; Kesson et al., 1994] demon- strate that the basaltic part of the subducted lithosphere can sink into the lower mantle. The seismic structure of the mantle

under subduction zones [Fukao et al., 1992] and numerical simulation [e.g., Tackley et al., 1993; Christensen and Hof- mann, 1994] have also demonstrated that the subducted slab can sink into the lower mantle. The location of large-scale up- welling from the lower mantle (superplumes), demonstrated by the three-dimensional seismic structure of the whole mantle

[Fukao, 1992], corresponds with the limited occurrence of the "pure" HIMU end-member in Polynesia in the South Pacific and St. Helena in the South Atlantic Ocean, although there are some other OIBs with HIMU-like signature, such as Canary Is- lands [Marcantonio eta/., 1995]. This suggests that the HIMU geochemical reservoir resides near the core-mantle boundary. The geochemical signature of HIMU basalts suggests the in- volvement of the subducted oceanic crust in the HIMU geo- chemical reservoir. It is thus suggested that the subducted oce- anic crust sinks into the lower mantle, incorporated in forming the HIMU geochemical reservoir in the lower mantle, and then ascends back to the Earth's surface as mantle plumes.

Conclusions

The chemical compositions of HIMU basalts are systemati- cally different in major and trace elements compared with non-

HIMUbasalts from Polynesian hotspots; HIMU primary mag- mas are more enriched in Fe203', MnO, and CaO but depleted in SiO2, K20, and P205 than non-HIMU primary magmas. This suggests that HIMU magmas originate from more fertile source material than non-HIMU magmas, indicating significant in- volvement of subducted oceanic crust in the genesis of the HIMU geochemical reservoir.

HIMU magmas are less enriched in highly incompatible trace elements than non-HIMU magmas, except for Nb. This feature can be partly explainedby involvement of sedimentary components in non-HIMU sources, but relative enrichment of Nb in HIMU requires some fractionation processes in the lower mantle. This is consistent with whole mantle scale recycling of the subducted oceanic crust, suggested by density relation- ships between the subducted lithosphere and the mantle and the seismic structure of the whole mantle.

Acknowledgments. Hayao Morinaga and his coworkers have kindly supplied us their samples from the Polynesian islands. Pb isotope analyses were obtained with great help from Susumu Nohda and Bor- ming Jahn. We also thank Kyoichi Ishizaka for help during chemical separation for Pb isotope analyses, Hironao Shinjo and Kei Hirose for helpful discussions, Shigeru Yamashita for providing us his unpublished data, and Henri Malarde and his family for their hospitality in Tubuai. We thank William White, Erik Hauri, and an anonymous referee for critical reviews. This work was supported by Research Fellowships of the Japan Society for the Promotion of Science for Young Scientists to T.K., grants-in-aid (05231106 and 08404029) from the Ministry of Edu- cation, Science and Culture of Japan, and a Special Coordination Fund for Promoting Science and Technology (Superplume) from Science and Technology Agency of Japan to Y.T.

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H. G. Barsczus, G6ofluides-Bassins-Eau, (UMR 5569 - CNRS et UM2), Insfitut des Sciences de la Terre, de l'Eau et de l'Espace de Montplellier, Universit6 de Montpellier II, 34095 Montpellier Cedex 5, France (e-mail: barsczus@dstu.univ-montp2.fr)

T. Kogiso, Department of Earth and Planetary Sciences, Tokyo Insti- me of Technology, Ookayama 2-12-1, Meguro, Tokyo 152, Japan. (e- mail: kogiso @ geo.titech.ac.jp)

G. Shimoda, Graduate School of Human and Environmental Studies, Kyoto University, Yoshida-Nihonmatsu, Kyoto 606-01, Japan. (e-mail: shimoda@gaia.h.kyoto-u.ac.jp)

Y. Tatsumi, School of Earth Sciences, Faculty of Integrated Human Studies, Kyoto University, Yoshida-Nihonmatsu, Kyoto 606-01, Japan. (e-mail: tatsumi@gaia. h.kyoto-u.ac.jp)

(Received January 3, 1996; revised December 5, 1996; accepted December 11, 1996)