evaluation of the oceanic crust-recycling model for himu oib

The Petrology and Geochemistry of St. HelenaAlkali Basalts: Evaluation of the OceanicCrust-recycling Model for HIMUOIB

HIROSHI KAWABATA*,TAKESHI HANYU, QING CHANG,JUN-ICHI KIMURA, ALEXANDER R. L. NICHOLS ANDYOSHIYUKI TATSUMIINSTITUTE FOR RESEARCH ON EARTH EVOLUTION (IFREE), JAPAN AGENCY FOR MARINE^EARTH SCIENCE

AND TECHNOLOGY (JAMSTEC), YOKOSUKA 237-0061, JAPAN

RECEIVED SEPTEMBER 15, 2009; ACCEPTEDJANUARY 20, 2011

We present major element, trace element, and petrographic data on

alkali basalts from St. Helena, and examine the geochemical charac-

teristics of a recycled component involved in the source of HIMU

(206Pb/ 204Pb420·5) ocean island basalts. Petrographic and com-

positional variations in the St. Helena basalts are best explained by

the combined effect of fractional crystallization and accumulation

of phenocrysts. Primary melt compositions are estimated by correcting

for the effects of crystal^liquid differentiation by reconstructing the

order of crystallization and the relative amount of fractionated

phases.This calculation indicates that the St. Helena alkali basalts

are derived from a common primary magma with 14^20 wt %

MgO. Simple partial melting of fertile mantle peridotite, depleted

mid-ocean ridge basalt (MORB)-source mantle, or garnet pyroxen-

ite fails to produce the St. Helena primary melt. Instead, this

primary melt can be reproduced if there are contributions from

ancient recycled oceanic crust and depleted peridotite [(Rb/

Nb)PM¼ 0·38^0·80]. Subducted sediment can be excluded to

explain the low (Rb, Ba, U)/Nb and Ce/Pb of St. Helena basalts.

Geochemical modeling using major and trace element abundances,

together with Sr, Nd, Pb, and Hf isotope ratios, indicates that the

St. Helena primary melt can be formed by 1^2% melting of a perido-

titic source that was refertilized by a small amount (8^18%) of

melt derived from recycled oceanic crust. This source has a similar

trace element pattern to modern normal (N)-MORB, but element

abundances are 0·1^0·2 times N-MORB values. The calculated

recycled crust has a wide range of present-day Pb isotopic ratios

(206Pb/ 204Pb of 21·7^79·3 and 208Pb/ 204Pb of 40·8^89·3),87Sr/ 86Sr of 0·7018^0·7028, 143Nd/144Nd of 0·51274^0·51285, and176Hf/177Hf of 0·28262^0·28293 after a residence time of 1·2^2·8Gyr. Rb, Ba, Pb, Sr, and light rare earth element abundances in the

recycled crust are depleted compared with modern N-MORB, where-

asTh, U, Sm, and Nd abundances fall within the range of compos-

itional variations in modern N-MORB. The trace element

compositions of the recycled oceanic crust can be explained by element

behavior during seafloor alteration and subduction zone dehydration

of oceanic crust. Therefore, recycling of ancient subducted oceanic

crust is a potential process for producing the St. Helena HIMU

basalts.

KEY WORDS: geochemistry; HIMU OIB; melting model; recycled

oceanic crust; St. Helena

I NTRODUCTIONTo explain the diversity in the Sr^Nd^Pb isotope compos-itions of mid-ocean ridge basalts (MORB) and oceanisland basalts (OIB), several discrete mantle componentssuch as HIMU, EM1, and EM2 have been proposed toexist in the Earth’s mantle (e.g. Zindler et al., 1982; White,1985; Zindler & Hart, 1986), in addition to depletedMORB-source mantle (DMM) and ubiquitous mantle

*Corresponding author.Telephone:þ81-46-867-9808.Fax:þ81-46-867-9808. E-mail: [email protected]

� The Author 2011. Published by Oxford University Press. Allrights reserved. For Permissions, please e-mail: [email protected]

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components called FOZO or C (Galer & O’Nions, 1985;Hart et al., 1992; Stracke et al., 2005). Heterogeneity in themantle chemistry may be caused by subducted oceaniccrust and sediment (Hofmann & White, 1982; Weaver,1991; Chauvel et al., 1992; Beier et al., 2007), metasomatizedlithospheric mantle (Halliday et al., 1995; Niu & O’ Hara,2003; Workman et al., 2004; Pilet et al., 2005), and delami-nated continental lower crust (White & Duncan, 1996;Tatsumi, 2000;Willbold & Stracke, 2006).One of the classical hypotheses assumed that ancient

subducted oceanic crust played an important role in pro-ducing OIB with a HIMU signature that is defined by ex-tremely radiogenic Pb isotopic ratios (206Pb/ 204Pb420·5;Hart, 1988). This hypothesis is commonly favored becausedehydration and seafloor alteration are expected to pro-duce the parent/daughter (P/D) ratios required to createthe isotopic characteristics of the HIMU OIB source(Zindler & Hart, 1986; Brenan et al., 1995; Keppler, 1996;Weaver et al., 1996; Hanyu & Kaneoka, 1997; Kogiso et al.,1997; Hilton et al., 2000; Stracke et al., 2003).Two approaches have been taken to examine quantita-

tively the oceanic crust-recycling hypothesis. One is a for-ward approach, in which the compositions of ancientsubducted oceanic crust are input considering element mo-bility during alteration and dehydration (e.g. Chauvelet al., 1992; Weaver et al., 1996; Kogiso et al., 1997). Then,petrogenetic conditions are explored to explain the geo-chemistry of HIMU OIB, although the modeling resultslargely depend on the initial subducted slab compositionsthat are used. For example, Chauvel et al. (1992) concludedthat trace element abundances and isotopic ratios ofHIMUOIB fromTubuai can be explained by partial melt-ing of a homogeneous mixture composed of peridotite andancient subducted oceanic crust. However, this conclusionis not the case for the dehydrated slab compositions esti-mated by Kogiso et al. (1997). They pointed out the import-ance of element fractionation in the presence of perovskiteunder lower mantle conditions. The discrepancy in model-ing results between the two studies is mainly caused bythe difference in the initial subducted slab compositionsused in the model. Therefore, accurate estimation of thesubducted slab composition is essential in the forward ap-proach; such estimation is not easy. A number of experi-ments have revealed element mobility during slabdehydration at specific P^T^H2O conditions (e.g. Brenanet al., 1995; Kogiso et al., 1997; Kessel et al., 2005). In add-ition, the extent of alteration is heterogeneous within theoceanic crust (Staudigel et al., 1996; Alt & Teagle, 2003;Bach et al., 2003; Kelley et al., 2003), and chemical modifi-cation during slab subduction strongly depends on theslab P^T trajectory.The above difficulty may be overcome by an inverse ap-

proach. This approach does not require an initial sub-ducted slab composition to be assumed, because the

composition is the outcome of the calculation. Using aninverse approach, previous studies constrained the P/Dratios in the subducted slab required to explain theradiogenic isotopic ratios of HIMU OIB, as a function ofthe initial isotopic ratios and the residence time of the sub-ducted slab in the mantle (Hauri & Hart, 1993; Strackeet al., 2003). These studies implicitly assumed that the iso-topic ratios of the time-evolved subducted slab wereequivalent to those of HIMU OIB as a first-order estima-tion. This is unreasonable, however, because the recycledoceanic crust alone is not adequate to be the source ofOIB (Hofmann & White, 1982; Niu & O’Hara, 2003;Prytulak & Elliott, 2007). Thus previous conclusionsshould be re-examined using a more appropriate petrogen-etic model.This study mainly addresses two issues. First, we esti-

mate the primary magma compositions based on petro-graphic, major and trace element data for HIMU alkalibasalts from St. Helena in the South Atlantic Ocean.Second, we perform geochemical modeling using theestimated primary melt compositions and published iso-topic ratios of St. Helena basalts. Our primary goal is toanswer the question of what chemical composition isrequired for the recycled ancient oceanic crust if theoceanic crust-recycling hypothesis is correct. To answerthis question, we expand upon previous inverse models,assuming that two components, peridotite and subductedoceanic crust, contributed to the genesis of St. Helenabasalts.

GEOLOGICAL BACKGROUNDSt. Helena is situated on 39 Ma oceanic lithosphere(Cande & Kent, 1992), c. 750 km east of the Mid-AtlanticRidge. The island reaches a height of 823m above sealevel and has a surface area of c. 122 km2. The subaerialpart of the island (c. 60 km3), which is about 5% of itstotal volume (Baker, 1969), consists of alkaline rocks ran-ging from highly porphyritic alkali basalts (e.g. ankara-mite) to trachyte and phonolite. The subaerial portion ofSt. Helena is dominated by a pile of basaltic lava flowserupted from two main centers (an older NE volcano anda younger SW volcano), associated with subordinate pyro-clastic rocks and trachytic to phonolitic intrusions. Thestratigraphic basal sequence of the NE volcano is exposedaround Prosperous Bay (Fig. 1a). This sequence consists ofvolcaniclastic rocks intruded by basaltic dyke swarms.Unconformably overlying the basal volcaniclastic rocksare shield-forming basaltic lava flows and pyroclastic de-posits (Baker, 1968a, 1969). Although aphyric to sparselyphyric basalts are common in the lava flows, highly por-phyritic alkali basalts are also found.Baker (1968a, 1969) classified the succession of the SW

volcano into three units based on erosional unconformities.The Lower Shield, which is made up of basaltic lava flows

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and pyroclastic deposits, is overlain by a thick pile of bas-altic lava flows interbedded with pyroclastic deposits, con-stituting the Main Shield. Younger basaltic andtrachy-andesitic lava flows defined as the Upper Shieldpartly cover the Main Shield of the NE volcano aroundProsperous Bay Plain (Fig. 1a). These shield-forming rocksin the SW volcano were intruded by trachytic and phonoli-tic intrusions during the final stage of volcanic activity.

K^Ar ages of 7^14 Ma have been reported for thevolcanic rocks that make up the NE and SW volcanoes(Abdel-Monem & Gast, 1967; Baker et al., 1967), butChaffey et al. (1989) suggested a more restricted rangeof 7^9 Ma. Of the subaerial part of St. Helena 95% ofthe volume is estimated to be alkali basalt, 4%trachyandesite, and 1% trachytes to phonolites (Baker,1968b).

Fig. 1. Geological map of St. Helena (modified from Baker,1969). In (a) open circles represent sample localities labeled with sample number. In(b), the age of the rocks from the Cameroon volcanic line and dredge samples from seamounts are from the compilations of Marzoli et al.(2000) and O’Connor & le Roex (1992). (b) is modified from O’Connor & le Roex (1992).

KAWABATA et al. ST. HELENA ALKALI BASALTS

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Seismic tomography has revealed the presence of a low-VP anomaly beneath St. Helena to c. 650 km depth, whichmerges with a deeper anomaly under Ascension at c.1000 km depth (Montelli et al., 2004).This seismic anomalyis recognized as evidence for a deep root to the St. Helenahotspot. From St. Helena to the west coast of Africa, abroad volcanic trail extends for over 2000 km (Fig. 1b).40Ar/ 39Ar ages of this seamount chain become progressive-ly older with increasing distance from St. Helena to a max-imum of 80^82 Ma (O’Connor & le Roex, 1992).Therefore, the plume that resulted in St. Helena basalts ofMiocene age has been active for at least the past 82 Myr.Part of the volcanic trail from St. Helena to the west coastof Africa is overprinted by igneous activity associatedwith the Cameroon line, a Y-shaped chain of Tertiary toQuaternary intra-plate volcanoes that extend across theoceanic to continental boundary centering around theactive volcano Mt. Cameroon (Fitton & Dunlop, 1985;Marzoli et al., 2000), and an older plutonic ring complexof 66^30 Ma (Cantagrel et al., 1978; Jacquemin et al., 1982).Plate reconstruction studies (O’Connor & Duncan, 1990;Lawver et al., 1999) suggest that the St. Helena hotspotwas originally located at a triple junction during the initialbreak-up of the Gondwana supercontinent (Sears et al.,2005).

ANALYT ICAL METHODSPhenocryst abundances were determined by point count-ing using more than 1500 points per thin section.Phenocrysts were usually defined as crystals with longaxes40·3mm.Vesicles, which were often observed in lavaflow samples, were excluded from the total number ofcounted points. Table 1 lists representative modal abun-dances in the St. Helena basalts.Mineral chemistries were determined using an electron-

probe microanalyzer (JEOLJXA-8800) at JAMSTEC. Anaccelerating voltage of 15 kV, beam current of 15 nA, andcounting time of 20 s for peaks and 10 s for backgroundswere used to analyze pyroxene and plagioclase. For olivine,the counting time was set to 100 s for Mn, Ca, and Ni,and 20 s for other elements at 20 kVand 25 nA (Shukuno,2003). The raw signals were corrected using a ZAF correc-tion procedure (e.g. Goldstein et al., 2003). Representativemineral compositions in St. Helena basalt lavas are listedin Supplementary Data AppendixTables A1^A4 (availableat http://www.petrology.oxfordjournals.org).For bulk-rock chemical analyses, all samples were pul-

verized in an alumina vibration mill after initially beingbroken up using a rock hammer. For the samples collectedfrom near the coastline, desalinization was performedusing tap water before washing in purified water. Rockpowder (0·4 g, dried at 1108C) was mixed with 4 g ofLi2B4O7 flux and fused in a Pt crucible for analysis ofmajor elements. Major element analyses were performed

by X-ray fluorescence spectrometry (XRF) using aRigaku Simultix 12 at JAMSTEC, following the procedureof Tani et al. (2006). Trace elements were determined byXRF on pressed powder pellets with a Rigaku RIX3000at JAMSTEC. The spectrum overlaps of TiKa line toBaLa line, YKb line to NbKb line, SrKb line to ZrKaline, and RbKb line toYKa line were corrected followingthe procedures of Goto & Tatsumi (1996). Matrix effectswere corrected by peak over background methods pro-posed by Champion et al. (1966). All discussion in thispaper refers to analyses that have been normalized to100% on a volatile-free basis with total iron calculated asFeO. Representative major and trace element compositionsfor the lavas are reported inTable 1.Trace elements were also analyzed by inductively

coupled plasma mass spectrometry (ICP-MS) using anAgilent 7500ce system fitted with PFA sample introductionand a Pt-inject torch system.The ICP-MS systemwas oper-ated with no collision gas and in multi-tune acquisitionmode. This combination allowed a wide range of elementsto be precisely determined using a pulse counting detector,and also a hydrofluoric acid sample solution to be de-livered directly into the plasma. Sample dissolution, prep-aration and measurement have been described by Changet al. (2003). Accurately weighed 100mg aliquots of pow-dered samples were digested in screw-cap PFA vials withmixed acids (0·65ml HClO4 and 2ml HF) on a hot plateat 130^1408C for 3 days. After drying, 1ml HClO4 wasadded, and the vials were heated at 1608C overnight.They were then opened to evaporate the solutions at agradually increased temperature of up to 1908C to driveoff excess Hf and convert the fluoride into chlorides. Theresidues were taken into 2ml of 6mol l�1 HNO3, moder-ately heated and then evaporated at 1208C to incipient dry-ness. The final dissolution was performed with 10ml of2% HNO3. Prior to analysis, In and Bi were added to ali-quots of sample solutions as internal standards to correctfor drift during measurements. Measured Eu and Gd con-centrations were corrected for oxide and hydroxide inter-ferences. The analytical results of repeated analyses ofwell-established reference standards (JB-2 and BHVO-2)listed in Table 1 agreed very well with the recommendedvalues (Imai et al.,1995; Jochum et al., 2005), demonstratingthat the trace elements were determined to a high accur-acy and reproducibility.

PETROGRAPHYRock samples were collected from all stratigraphic unitsfound above sea level (Fig. 1). Petrographical features,including phenocryst assemblage, mineral abundances,and phenocryst size, roughly correlate with bulk-rockcompositions. The phenocryst assemblage changes fromolivine^clinopyroxene dominant to plagioclase^olivinedominant with decreasing MgO (Fig. 2). In addition, the


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Table 1: Mineral abundances, major and trace element compositions of volcanic rocks from St. Helena

Unit: NE volcano

Latitude: �15·928 �15·928 �15·928 �15.941 �15·941 �15·946 �15·906 �15·907 �15·907 �15·911

Longitude: �5·651 �5·657 �5·655 �5·646 �5·646 �5·648 �5·699 �5·701 �5·700 �5·700

Rock type: Basalt Basalt Picro B Trachyte Trachyte Basanite TB TB BTA Basalt

Sample no.: SH-32 SH-34 SH-35 SH-44A SH-44B SH-45 SH-57 SH-58 SH-59 SH-60

XRF (wt %)

SiO2 46·22 44·57 43·52 61·41 62·38 44·35 45·63 46·00 51·24 44·98

TiO2 3·19 3·70 1·99 0·28 0·28 2·65 3·81 3·51 1·62 3·78

Al2O3 16·27 15·25 9·73 18·12 18·86 12·54 16·45 16·34 17·91 15·29

Fe2O3 13·00 13·53 13·10 6·00 3·97 12·59 12·85 13·03 10·66 13·61

MgO 4·97 4·70 15·72 0·30 0·22 11·06 4·43 4·97 2·38 4·94

CaO 10·11 10·38 12·24 0·91 1·17 12·43 10·15 8·86 5·36 11·82

MnO 0·18 0·18 0·18 0·07 0·05 0·17 0·18 0·19 0·20 0·16

Na2O 3·33 3·47 1·51 6·82 7·05 2·30 3·53 3·70 5·97 2·55

K2O 1·03 1·22 0·52 4·67 4·70 0·81 1·52 1·72 2·65 0·98

P2O5 0·52 0·58 0·28 0·07 0·08 0·40 0·87 0·78 0·59 0·56

Total 98·81 97·58 98·79 98·65 98·76 99·29 99·42 99·09 98·57 98·66

XRF (ppm)

Ba 852 866 343 233

Ni 2 6 30 56

Cu 62 89 46 84

Zn 106 109 83 130 120 90 109 102 134 106

Pb 9 6 3

Th 21 24 5 4

Rb 110 96 26 16

Sr 59 73 874 709

Y 61 84 27 22

Zr 1020 1158 290 201

Nb 208 193 64 43

ICP-MS (ppm)

Sc 18·4 34·0 33·2 14·4

Co 38·8 71·1 59·2 34·0 13·4

Ni 62·0 376 227 29·4

Cu 88·4 59·5 82·3 33·1 19·2

Rb 18·0 10·3 15·8 28·8 55·8

Sr 559 319 455 795 763

Y 27·3 16·4 22·1 31·6 31·5

Zr 231 135 194 311 541

Nb 44·9 25·4 37·9 69·9 109

Cs 0·213 0·280 0·170 0·388 0·516

Ba 258 147 217 413 616

La 33·6 20·1 28·8 51·5 79·9

Ce 71·6 42·9 60·8 107 154

Pr 8·88 5·38 7·58 12·9 16·7

Nd 37·0 22·7 31·3 53·0 60·3

Sm 8·12 5·01 6·73 10·5 10·5

Eu 2·61 1·62 2·16 3·33 3·22

Gd 7·74 4·90 6·34 9·55 8·59

Tb 1·17 0·724 0·958 1·41 1·32

Dy 6·20 3·87 5·09 7·20 6·74

Ho 1·14 0·709 0·931 1·32 1·25

Er 3·01 1·88 2·45 3·48 3·46

Tm 0·390 0·241 0·314 0·444 0·461

Yb 2·33 1·44 1·93 2·67 2·95

Lu 0·337 0·212 0·271 0·388 0·431

Hf 5·88 3·71 5·07 7·38 10·9

Ta 2·84 1·65 2·46 4·36 6·85

Tl 0·027 0·007 0·051 0·038 0·055

Pb 1·78 0·938 1·38 2·47 5·04

Th 3·60 2·17 3·33 5·41 10·9

U 1·02 0·630 0·951 1·48 2·81

Mode (vol. %)

groundmass 61 99 49 72 90 94 99 88

olivine 3 Tr. 19 11 2 2 Tr. 2

clinopyroxene 13 Tr. 31 17 Tr. 2 Tr. 2

feldspar 21 1 2 0 8 3 Tr. 7

amphibole 0 0 0 0 0 0 Tr. 0

Fe–Ti oxide 1 Tr. Tr. Tr. Tr. Tr. Tr. Tr.

(continued)


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Table 1: Continued

Unit: NE volcano SW volcano, Lower Shield

Latitude: �15·909 �15·909 �15·911 �16·009 �16·014 �16·013 �16·009 �16·003 �16·003 �16·003

Longitude: �5·706 �5·706 �5·705 �5·725 �5·727 �5·726 �5·725 �5·714 �5·714 �5·715

Rock type: TB TB Basalt Basalt TB Basanite Basalt Basalt TA Basalt

Sample no.: SH-62 SH-64 SH-65 SH-21 SH-23 SH-24 SH-25 SH-26 SH-27 SH-28

XRF (wt %)

SiO2 45·98 47·04 44·36 45·16 49·31 43·43 44·95 45·59 53·95 45·97

TiO2 3·13 2·95 3·61 2·10 2·11 2·65 2·31 3·59 0·77 3·94

Al2O3 16·75 16·60 15·95 10·54 17·06 10·95 11·07 15·19 17·96 16·58

Fe2O3 12·53 12·95 13·40 12·75 11·78 13·03 13·19 13·55 10·30 13·82

MgO 4·09 4·04 4·73 14·55 2·24 11·61 13·99 6·52 1·17 3·73

CaO 9·04 8·72 11·18 11·67 6·82 11·98 10·76 10·01 3·90 8·86

MnO 0·19 0·20 0·16 0·17 0·22 0·18 0·18 0·19 0·25 0·15

Na2O 4·34 3·97 3·35 1·59 5·46 2·30 2·30 3·39 6·70 3·44

K2O 1·69 1·75 1·41 0·62 2·16 0·83 0·71 1·19 2·86 1·43

P2O5 1·12 1·15 0·76 0·29 0·95 0·40 0·36 0·56 0·39 0·60

Total 98·84 99·36 98·90 99·45 98·13 97·36 99·82 99·76 98·24 98·52

XRF (ppm)

Ba 409 339 512 275 674 325

Ni 15 43 85 4 32

Cu 39 67 13 65 7 34

Zn 110 108 102 86 125 99 92 112 133 99

Pb 3 2 4 2 6 2

Th 7 5 8 4 14 5

Rb 31 48 49 23 68 24

Sr 843 807 751 602 690 693

Y 32 26 35 27 31 28

Zr 352 276 438 249 607 279

Nb 79 63 95 53 125 60

ICP-MS (ppm)

Sc 12·9 33·8 27·6 28·5

Co 30·2 64·8 61·6 65·8

Ni 14·5 300 289 301

Cu 40·5 74·0 146 58·7

Rb 35·1 13·1 16·2 12·9

Sr 844 385 459 352

Y 31·5 18·2 20·2 19·2

Zr 329 178 189 168

Nb 76·9 35·0 35·8 31·3

Cs 0·368 0·119 0·149 0·100

Ba 432 196 225 176

La 63·8 24·3 27·6 24·0

Ce 130 51·2 59·0 51·3

Pr 15·2 6·21 7·29 6·38

Nd 59·5 25·7 30·5 26·6

Sm 11·1 5·42 6·49 5·76

Eu 3·44 1·77 2·09 1·86

Gd 9·65 5·30 6·17 5·59

Tb 1·41 0·792 0·913 0·838

Dy 7·15 4·20 4·79 4·44

Ho 1·29 0·767 0·852 0·821

Er 3·41 2·04 2·21 2·16

Tm 0·428 0·262 0·275 0·276

Yb 2·59 1·57 1·63 1·71

Lu 0·382 0·225 0·231 0·237

Hf 7·82 4·41 4·84 4·29

Ta 4·86 2·23 2·29 2·02

Tl 0·037 0·015 0·046 0·008

Pb 3·06 1·32 1·42 1·16

Th 6·87 3·23 2·93 2·56

U 1·26 0·896 0·877 0·840

Mode (vol. %)

groundmass 97 91 90 66 82 77 83 90

olivine Tr. 2 2 25 13 10 2 5

clinopyroxene 0 Tr. 2 9 5 13 6 Tr.

feldspar 3 6 6 0 0 0 9 5

amphibole 0 0 0 0 0 0 0 0

Fe–Ti oxide Tr. Tr. Tr. Tr. Tr. Tr. Tr. 1

(continued)


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Table 1: Continued

Unit: SW volcano, Lower Shield SW volcano, Main Shield

Latitude: �16·007 �16·007 �16·007 �16·007 �16·007 �15·927 �15·927 �15·928 �15·929 �15·929

Longitude: �5·716 �5·716 �5·717 �5·718 �5·719 �5·720 �5·720 �5·719 �5·718 �5·718

Rock type: Basalt Basalt TB TB Basalt TB TB TB BTA Basalt


XRF (wt %)

SiO2 44·78 45·65 47·06 47·10 44·67 50·78 51·20 49·21 51·84 44·97

TiO2 3·25 3·16 3·13 3·13 2·66 1·83 1·73 2·10 1·42 3·62

Al2O3 15·53 16·25 17·15 17·09 12·50 16·88 17·04 16·41 17·35 16·26

Fe2O3 12·46 12·63 12·83 12·77 13·34 12·72 12·74 13·80 11·73 13·67

MgO 6·20 5·63 3·74 4·03 11·77 2·73 2·54 2·61 2·68 5·45

CaO 12·18 11·29 8·14 8·09 11·12 6·10 5·82 6·48 5·05 10·52

MnO 0·19 0·17 0·17 0·23 0·19 0·27 0·32 0·31 0·25 0·18

Na2O 2·61 2·95 3·90 4·04 2·18 4·94 5·01 4·82 5·46 2·93

K2O 0·97 1·05 1·82 1·73 0·81 2·26 2·37 2·14 2·57 1·25

P2O5 0·50 0·53 0·81 0·81 0·41 1·03 0·98 1·25 0·72 0·62

Total 98·66 99·31 98·76 99·00 99·64 99·54 99·75 99·12 99·06 99·47

XRF (ppm)

Ba 245 261 432 429 532 512 484 594 322

Ni 82 54 11 3 4 7 13 4 46

Cu 48 57 17 18 18 18 10 10 52

Zn 90 96 94 98 93 133 138 121 121 103

Pb 2 3 3 3 4 5 3

Th 4 4 5 7 10 11 10 11 5

Rb 18 20 30 30 50 53 42 81 22

Sr 621 602 860 858 751 741 883 805 694

Y 23 25 29 29 33 33 35 32 26

Zr 215 220 307 304 469 490 402 503 250

Nb 47 50 71 71 97 101 90 106 59

ICP-MS (ppm)

Sc 28·0

Co 61·2

Ni 219

Cu 56·6

Rb 15·6

Sr 469

Y 21·4

Zr 192

Nb 39·1

Cs 0·216

Ba 220

La 29·6

Ce 62·2

Pr 7·58

Nd 30·9

Sm 6·49

Eu 2·09

Gd 6·11

Tb 0·92

Dy 4·84

Ho 0·884

Er 2·36

Tm 0·299

Yb 1·83

Lu 0·266

Hf 4·76

Ta 2·45

Tl 0·040

Pb 1·41

Th 3·22

U 0·913

Mode (vol. %)

groundmass 97 99 99 79 99 99 99 98 95

olivine 2 0 0 12 Tr. Tr. Tr. 1 2

clinopyroxene Tr. 0 0 9 Tr. Tr. Tr. Tr. Tr.

feldspar 0 0 0 0 Tr. Tr. Tr. Tr. 3

amphibole 0 Tr. Tr. 0 0 0 0 0 0

Fe–Ti oxide 0 Tr. Tr. Tr. Tr. Tr. 1 Tr. Tr.

(continued)


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Table 1: Continued

Unit: SW volcano, Main Shield

Latitude: �15·930 �15·929 �15·930 �15·930 �15·930 �15·931 �15·931 �15·931 �15·931 �15·932

Longitude: �5·717 �5·717 �5·717 �5·716 �5·716 �5·715 �5·715 �5·714 �5·714 �5·714

Rock type: Basalt TB Basalt Basalt Basalt Basalt Basalt Basalt Basalt TB


XRF (wt %)

SiO2 46·08 49·07 45·22 44·49 45·39 46·13 46·82 45·71 45·37 46·78

TiO2 2·73 2·71 2·90 3·09 3·86 3·67 3·21 2·57 3·55 3·79

Al2O3 14·82 16·88 14·50 14·59 15·77 16·21 15·80 12·33 15·52 16·14

Fe2O3 13·24 12·54 13·23 13·41 14·03 13·63 13·01 13·45 13·55 13·78

MgO 8·10 3·44 8·25 8·19 5·16 5·28 5·61 9·65 5·84 3·72

CaO 9·60 7·42 10·85 11·39 9·89 9·38 9·47 11·39 10·49 8·52

MnO 0·18 0·22 0·17 0·17 0·19 0·19 0·19 0·20 0·18 0·20

Na2O 2·92 4·45 2·55 2·38 3·46 3·57 3·42 2·01 2·72 3·69

K2O 0·88 1·96 0·97 0·90 1·24 1·28 1·38 0·79 1·09 1·42

P2O5 0·47 0·95 0·48 0·46 0·59 0·63 0·66 0·42 0·60 0·68

Total 99·00 99·64 99·13 99·07 99·55 99·95 99·56 98·52 98·90 98·71

XRF (ppm)

Ba 211 431 327 314 291

Ni 152 9 19 57 83

Cu 53 14 42 62 55

Zn 93 135 94 91 109 95 101 90 102 117

Pb 2 3 3 2 3

Th 4 7 5 5 4

Rb 15 38 24 25 15

Sr 499 731 655 639 659

Y 22 47 28 27 25

Zr 194 382 270 284 238

Nb 39 86 58 63 52

ICP-MS (ppm)

Sc 24·5 27·1 19·9 31·2 15·3

Co 50·2 51·2 42·1 68·7 32·7

Ni 140 120 33·3 314 4·89

Cu 68·3 51·5 38·0 77·0 19·5

Rb 15·9 14·5 21·3 12·8 23·7

Sr 539 568 603 421 660

Y 22·9 23·4 28·8 21·3 32·3

Zr 206 210 273 187 305

Nb 42·7 41·9 55·8 37·3 64·0

Cs 0·148 0·145 0·197 0·146 0·297

Ba 261 254 321 222 512

La 32·4 31·8 41·8 28·5 47·4

Ce 67·1 66·8 86·1 59·6 97·3

Pr 8·16 8·29 10·4 7·37 11·9

Nd 33·3 34·3 42·3 30·6 48·5

Sm 7·01 7·28 8·76 6·49 9·93

Eu 2·30 2·34 2·80 2·06 3·17

Gd 6·64 6·81 8·18 6·17 9·36

Tb 1·00 1·02 1·22 0·921 1·37

Dy 5·28 5·35 6·47 4·87 7·21

Ho 0·960 0·987 1·19 0·894 1·32

Er 2·55 2·58 3·14 2·36 3·54

Tm 0·327 0·332 0·405 0·300 0·454

Yb 1·98 2·01 2·45 1·81 2·74

Lu 0·283 0·287 0·350 0·259 0·401

Hf 5·23 5·36 6·66 4·73 7·27

Ta 2·77 2·70 3·60 2·37 4·05

Tl 0·010 0·010 0·018 0·010 0·012

Pb 1·73 1·83 2·32 1·45 2·39

Th 3·76 3·58 5·03 3·15 5·21

U 0·977 0·948 1·13 0·807 1·32

Mode (vol. %)

groundmass 93 99 84 86 98 98 92 80 90 95

olivine 4 Tr. 9 6 Tr. Tr. 7 12 10 4

clinopyroxene 2 Tr. 7 7 Tr. Tr. Tr. 8 Tr. Tr.

feldspar Tr. Tr. 0 2 1 1 2 0 Tr. Tr.

amphibole 0 0 0 0 0 0 0 0 0 0

Fe–Ti oxide Tr. Tr. Tr. Tr. Tr. Tr. Tr. Tr. Tr. 1

(continued)


798

Dow



Table 1: Continued


Latitude: �15·932 �15·932 �16·001 �16·008 �16·009 �15·998 �15·999 �15·997 �16·003

Longitude: �5·714 �5·713 �5·747 �5·743 �5·743 �5·761 �5·761 �5·783 �5·716

Rock type: BTA TB Basalt Basalt Basalt TB BTA TB TA Basalt


XRF (wt %)

SiO2 51·63 50·42 45·12 44·52 45·11 49·74 52·26 47·13 56·18 45·33

TiO2 1·70 1·97 2·39 3·29 3·55 2·50 1·92 3·35 0·70 2·00

Al2O3 17·10 17·22 13·75 15·78 15·61 17·88 18·56 16·64 17·69 9·56

Fe2O3 12·69 12·39 13·03 13·52 13·13 10·35 9·62 13·10 10·17 11·71

MgO 2·10 2·57 10·77 7·35 5·63 3·14 2·05 4·40 0·58 15·42

CaO 5·43 6·36 11·32 10·92 10·09 7·01 5·47 8·33 2·74 12·45

MnO 0·23 0·23 0·18 0·17 0·18 0·19 0·20 0·25 0·33 0·16

Na2O 5·02 4·77 2·22 2·61 3·20 5·27 5·21 3·84 6·19 1·19

K2O 2·25 2·11 0·69 0·86 1·46 2·15 2·47 1·73 3·46 0·48

P2O5 0·98 1·00 0·36 0·45 0·62 0·92 0·64 0·79 0·25 0·29

Total 99·13 99·02 99·83 99·46 98·58 99·14 98·39 99·55 98·28 98·58

XRF (ppm)

Ba 551 527 205 316 725 393 856

Ni 2 4 87 48 11 6

Cu 19 18 64 40 6 24

Zn 130 129 83 94 97 119 125 109 140 77

Pb 5 3 2 3 7 2 5

Th 10 9 4 6 15 6 14

Rb 45 49 19 32 62 29 71

Sr 714 892 679 687 770 754 585

Y 32 34 22 27 39 33 46

Zr 485 454 182 259 548 307 658

Nb 100 99 39 58 124 72 144

ICP-MS (ppm)

Sc 25·7 6·69 37·2

Co 66·9 18·5 62·2

Ni 224 339

Cu 62·4 14·0 73·1

Rb 13·1 48·7 8·44

Sr 487 784 277

Y 22·2 35·9 16·5

Zr 167 464 135

Nb 33·1 102 26·4

Cs 0·105 0·510 0·061

Ba 211 561 143

La 26·9 70·5 20·5

Ce 53·8 140 43·4

Pr 6·74 15·8 5·36

Nd 28·1 60·6 22·5

Sm 5·93 11·4 4·93

Eu 1·99 3·51 1·59

Gd 5·94 10·1 4·78

Tb 0·869 1·49 0·72

Dy 4·62 7·84 3·83

Ho 0·874 1·44 0·702

Er 2·32 3·90 1·84

Tm 0·292 0·510 0·238

Yb 1·73 3·16 1·41

Lu 0·248 0·460 0·206

Hf 4·31 9·72 3·56

Ta 2·11 6·4 1·65

Tl 0·018 0·125 0·005

Pb 1·69 4·98 0·82

Th 2·76 10·9 2·27

U 0·779 3·21 0·625

Mode (vol. %)

groundmass 98 99 73 78 94 98 94 100 53

olivine 1 Tr. 10 4 2 Tr. Tr. 0 15

clinopyroxene Tr. Tr. 9 1 1 Tr. Tr. 0 33

feldspar Tr. Tr. 8 16 3 1 3 0 0

amphibole 0 0 0 0 0 Tr. 2 0 0

Fe–Ti oxide Tr. Tr. Tr. Tr. Tr. 1 1 0 Tr.

(continued)


799

Dow



Table 1: Continued


Latitude: �15·996 �16·003 �16·001 �15·997 �15·966 �15·918

Longitude: �5·700 �5·694 �5·696 �5·702 �5·745 �5·715

Rock type: TB TB TB Picro B TB TB Basalt TB TB TB


XRF (wt %)

SiO2 45·39 48·03 49·24 44·07 47·71 46·63 44·63 48·93 50·29 47·16

TiO2 3·41 2·82 2·43 2·55 2·83 3·12 3·35 2·71 1·80 3·20

Al2O3 15·29 17·20 16·91 12·36 17·54 17·32 15·35 16·80 16·73 16·73

Fe2O3 14·26 13·30 13·26 13·83 13·73 12·02 13·50 12·50 12·50 13·14

MgO 6·20 2·70 2·65 11·73 3·27 4·32 6·62 3·56 2·48 4·72

CaO 8·36 7·79 6·88 12·05 7·09 9·16 10·76 7·48 5·99 8·04

MnO 0·20 0·16 0·20 0·22 0·18 0·18 0·18 0·22 0·25 0·20

Na2O 4·07 4·17 4·64 1·77 3·30 4·08 2·95 4·44 5·30 3·85

K2O 1·57 1·70 1·98 0·61 1·69 1·61 1·17 1·91 2·25 1·71

P2O5 0·73 1·14 1·32 0·34 0·91 0·81 0·63 0·97 1·03 0·91

Total 99·48 99·01 99·52 99·54 98·23 99·25 99·14 99·51 98·61 99·67

XRF (ppm)

Ba 414 497 667 464 378 479 500 411

Ni 72 36 11 30 29 2 7 14

Cu 53 25 20 29 30 11 18 24

Zn 102 123 144 94 132 104 98 125 144 115

Pb 3 5 4 2 3 3 5 2

Th 6 7 9 8 7 7 10 7

Rb 32 35 50 34 28 34 49 29

Sr 837 753 764 757 865 730 725 773

Y 28 49 35 44 29 36 33 30

Zr 321 339 408 377 305 379 470 332

Nb 71 78 93 86 68 85 99 78

ICP-MS (ppm)

Sc 30·8 21·1

Co 78·2 44·5

Ni 219 82·4

Cu 44·2 48·5

Rb 11·4 20·5

Sr 430 652

Y 18·8 25·3

Zr 159 243

Nb 30·6 54·3

Cs 0·073 0·195

Ba 192 317

La 24·3 41·0

Ce 51·5 84·2

Pr 6·36 10·1

Nd 26·8 40·6

Sm 5·79 8·16

Eu 1·89 2·61

Gd 5·53 7·55

Tb 0·83 1·11

Dy 4·42 5·85

Ho 0·801 1·07

Er 2·11 2·83

Tm 0·267 0·369

Yb 1·61 2·22

Lu 0·233 0·323

Hf 4·28 5·90

Ta 1·99 3·44

Tl 0·025 0·013

Pb 1·13 2·19

Th 2·71 4·72

U 0·741 0·993

Mode (vol. %)

groundmass 88 88 96 61 90 80 91 96 99 100

olivine 10 4 2 12 8 4 6 2 Tr. Tr.

clinopyroxene Tr. Tr. Tr. 20 Tr. Tr. Tr. Tr. Tr. Tr.

feldspar 1 7 2 6 Tr. 15 3 2 Tr. Tr.

amphibole 1 0 0 0 0 0 0 0 0 0

Fe–Ti oxide 1 Tr. 1 Tr. 1 Tr. Tr. Tr. Tr. Tr.

(continued)


800

Dow



Table 1: Continued

Unit: SW volcano, Upper Shield

Latitude: �15·931 �15·952 �15·947 �15·947 �15·952 �15·953 �15·952

Longitude: �5·658 �5·646 �5·648 �5·648 �5·650 �5·651 �5·650

Rock type: TA TB TB BTA Basalt Basalt Basalt TB TA TB


XRF (wt %)

SiO2 54·66 49·30 48·03 50·83 45·38 44·699 45·39 51·15 54·53 50·70

TiO2 0·79 2·77 2·68 1·92 3·57 3·187 3·34 1·91 0·77 2·01

Al2O3 18·00 17·97 16·60 17·01 15·97 14·747 16·40 17·33 17·85 17·11

Fe2O3 10·36 13·46 13·26 11·75 13·47 13·286 13·15 11·82 10·30 12·39

MgO 1·53 1·55 3·97 2·24 5·40 7·799 5·53 2·32 1·55 1·59

CaO 3·80 5·79 6·98 5·78 9·57 10·691 9·27 6·13 3·81 6·01

MnO 0·29 0·13 0·22 0·20 0·16 0·196 0·18 0·17 0·27 0·14

Na2O 5·93 4·50 4·11 4·95 3·00 2·419 3·20 4·93 6·09 4·87

K2O 2·96 1·91 1·69 2·19 1·24 1·072 1·40 2·17 2·92 2·27

P2O5 0·37 1·31 0·94 0·99 0·64 0·591 0·62 0·96 0·37 1·22

Total 98·69 98·69 98·47 97·87 98·39 98·69 98·46 98·89 98·45 98·32

XRF (ppm)

Ba 770 421 577 331 538 523

Ni 6 3 40 6 5

Cu 5 27 21 11 48 9 15

Zn 127 155 114 130 109 94 93 134 135 119

Pb 7 5 2 2 4 4

Th 15 7 10 6 10 9

Rb 72 34 41 25 46 47

Sr 692 729 649 867 783 874

Y 35 32 35 25 35 36

Zr 654 367 464 253 457 439

Nb 133 81 97 54 100 97

ICP-MS (ppm)

Sc 9·23 16·8 22·7

Co 16·3 38·8 45·2 5·49

Ni 5·29 41·5 94·4

Cu 27·0 46·5 68·7 8·12

Rb 40·4 22·7 18·9 67·2

Sr 796 627 600 661

Y 35·7 27·5 25·3 35·6

Zr 411 273 255 677

Nb 85·7 58·0 55·3 139

Cs 0·195 0·185 0·120 0·451

Ba 524 344 318 741

La 89·0 43·7 41·9 95·4

Ce 169 90·2 86·4 181

Pr 19·3 10·9 10·4 18·9

Nd 72·8 43·8 41·8 66·1

Sm 13·2 9·03 8·44 10·9

Eu 4·10 2·88 2·66 3·22

Gd 11·1 8·31 7·68 8·72

Tb 1·63 1·23 1·14 1·38

Dy 8·24 6·46 5·93 7·25

Ho 1·49 1·18 1·07 1·38

Er 3·97 3·11 2·85 3·98

Tm 0·514 0·396 0·366 0·555

Yb 3·15 2·42 2·21 3·64

Lu 0·465 0·34 0·316 0·54

Hf 9·08 6·63 6·21 13·7

Ta 5·32 3·73 3·55 8·39

Tl 0·023 0·045 0·014 0·060

Pb 3·66 2·18 2·25 6·19

Th 8·21 4·96 4·71 13·6

U 2·29 1·21 1·29 3·72

Mode (vol. %)

groundmass 94 93 94 99 98 95

olivine 2 3 6 Tr. Tr. 1

clinopyroxene Tr. 1 Tr. Tr. Tr. Tr.

feldspar 3 4 0 Tr. 1 4

amphibole 0 0 0 0 0 0

Fe–Ti oxide 1 Tr. Tr. 0 Tr. Tr.

(continued)


801

Dow



Table 1: Continued

Unit: SW volcano, Upper Shield SW volcano, late-stage intrusions

Latitude: �15·955 �15·970 �15·973 �15·973 �16·009 �15·985 �15·996

Longitude: �5·652 �5·653 �5·650 �5·649 �5·725 �5·716 �5·746

Rock type: TB TA TB TB Trachyte Trachyte Trachyte Trachyte Phonolite Trachyte


XRF (wt %)

SiO2 46·29 54·84 47·52 46·91 60·18 60·12 59·56 59·18 58·92 61·41

TiO2 3·44 0·73 2·89 2·93 0·03 0·05 0·23 0·27 0·07 0·11

Al2O3 16·35 18·04 16·23 15·83 18·71 18·40 17·94 17·99 19·58 18·23

Fe2O3 13·62 10·17 13·08 12·96 4·36 4·72 7·03 6·57 4·45 4·91

MgO 3·36 0·83 4·67 5·52 0·00 0·14 0·06 0·50 0·00 0·05

CaO 8·78 4·09 8·62 9·31 0·58 0·79 1·97 1·90 1·01 1·26

MnO 0·20 0·26 0·24 0·18 0·20 0·21 0·17 0·24 0·19 0·19

Na2O 3·56 5·99 3·79 3·79 8·47 8·38 7·36 7·04 8·99 7·62

K2O 1·54 3·32 1·50 1·40 5·02 5·00 4·30 4·69 4·94 5·01

P2O5 1·13 0·35 0·81 0·71 0·06 0·05 0·08 0·07 0·05 0·05

Total 98·28 98·62 99·34 99·53 97·60 97·85 98·70 98·44 98·20 98·83

XRF (ppm)

Ba 371 764 398 166 1335 306

Ni 40 46 n.d.

Cu 32 6 45 n.d. 2

Zn 140 130 112 112 209 247 166 155 194 191

Pb 4 6 4 15 8 16

Th 6 15 7 38 18 37

Rb 27 77 31 187 108 179

Sr 772 704 713 25 214 39

Y 34 34 29 58 63 35

Zr 283 653 328 1460 912 1141

Nb 68 136 72 200 222 128

ICP-MS (ppm)

Sc 16·9

Co 38·1 1·72 1·30 0·75

Ni 71·3

Cu 50·7 4·39 5·17

Rb 28·2 164 108 127

Sr 660 16·4 217 92·4

Y 27·5 69·1 44·0 55·4

Zr 306 1341 988 1135

Nb 62·9 166 198 194

Cs 0·238 1·91 1·15 1·72

Ba 376 125 1351 742

La 49·8 190 126 140

Ce 101 312 233 252

Pr 11·9 27·9 23·8 24·7

Nd 47·0 82·4 78·90 79·7

Sm 9·08 12·0 12·7 12·8

Eu 2·84 1·68 3·04 2·31

Gd 8·15 9·83 9·78 10·4

Tb 1·21 1·80 1·62 1·79

Dy 6·29 10·5 8·68 9·99

Ho 1·15 2·27 1·68 2·01

Er 3·06 7·10 4·99 6·01

Tm 0·397 1·06 0·709 0·856

Yb 2·42 6·92 4·76 5·65

Lu 0·353 1·03 0·710 0·833

Hf 6·94 27·9 20·4 23·74

Ta 3·95 14·4 11·6 13·02

Tl 0·029 0·283 0·143 0·167

Pb 2·55 15·8 9·71 10·8

Th 5·70 33·9 19·7 22·7

U 1·58 8·66 5·51 5·16

Mode (vol. %)

groundmass 91 98 93 87 86

olivine 2 2 7 0 Tr.

clinopyroxene Tr. Tr. Tr. Tr. 1

feldspar 6 Tr. Tr. 13 13

amphibole 0 0 0 0 0

Fe–Ti oxide Tr. Tr. Tr. Tr. Tr.

(continued)


802

Dow



Table 1: Continued

Unit: SW volcano, late-stage intrusions

Latitude: �15·982 �15·980 �15·980 �15·980 �15·979 �15·995 �16·002 �15·992 �15·968

Longitude: �5·667 �5·653 �5·657 �5·660 �5·666 �5·758 �5·717 �5·701 �5·653

Rock type: TA Trachyte Trachyte Trachyte BTA Trachyte TA Phonolite Trachyte TA TA

Sample no.: SH-67 SH-68 SH-69 SH-70 SH-71 SH-72 SH-79 SH-85 SH-91 SH-93 SH-94

XRF (wt %)

SiO2 54·57 57·41 56·67 56·39 52·63 61·99 55·94 59·47 60·10 54·08 54·23

TiO2 1·11 0·24 0·25 0·23 1·60 0·11 0·53 0·06 0·28 1·22 1·23

Al2O3 18·38 19·04 19·12 18·97 18·14 18·26 18·07 19·63 17·93 18·70 18·56

Fe2O3 10·97 7·67 7·97 7·89 12·09 5·01 9·59 4·36 6·84 11·12 11·03

MgO 0·24 0·24 0·24 0·56 1·12 0·03 0·55 0·00 0·11 0·81 0·78

CaO 3·17 2·18 2·53 2·48 3·96 1·12 3·16 1·21 2·23 2·97 3·08

MnO 0·06 0·38 0·32 0·33 0·15 0·21 0·29 0·19 0·22 0·13 0·13

Na2O 6·10 7·65 8·09 7·68 5·63 7·91 6·85 8·73 6·99 5·80 5·80

K2O 2·84 3·58 3·59 3·58 2·44 5·02 3·37 5·04 4·54 2·69 2·74

P2O5 0·89 0·15 0·16 0·14 1·21 0·05 0·24 0·06 0·09 0·99 1·01

Total 98·34 98·54 98·94 98·26 98·98 99·71 98·59 98·74 99·33 98·51 98·58

XRF (ppm)

Ba 832 699 749 666 814 227 1019 665

Ni 9 8 2 4

Cu 10 3 5 16 5 2 12

Zn 100 154 146 131 117 214 131 194 163 99 93

Pb 7 11 10 5 8 17 9 5

Th 17 25 22 13 16 36 20 15

Rb 76 113 106 63 86 179 110 72

Sr 751 502 522 877 570 33 177 777

Y 24 38 34 25 35 35 53 30

Zr 601 814 803 511 727 1165 980 630

Nb 122 135 130 106 139 132 202 121

ICP-MS (ppm)

Sc 3·25

Co 3·33 0·831 8·26

Ni 3·78

Cu 6·01 13·7

Rb 105 135 60·7

Sr 515 52·8 719

Y 35·6 58·5 26·3

Zr 816 1175 608

Nb 135 204 119

Cs 1·02 1·13 0·262

Ba 694 592 647

La 106 152 88·2

Ce 186 271 168

Pr 18·0 26·9 17·2

Nd 57·6 86·8 60·6

Sm 8·90 13·8 10·1

Eu 2·45 2·30 3·05

Gd 7·03 11·4 7·80

Tb 1·17 1·93 1·17

Dy 6·47 10·6 5·86

Ho 1·30 2·09 1·08

Er 3·99 6·20 2·92

Tm 0·60 0·896 0·396

Yb 4·09 5·91 2·53

Lu 0·617 0·852 0·362

Hf 16·1 25·0 12·3

Ta 9·80 13·6 7·62

Tl 0·06 0·053 0·038

Pb 10·6 11·4 5·73

Th 21·7 24·5 13·1

U 5·97 6·85 3·63

Mode (vol. %)

groundmass 99 81 84

olivine Tr. 0 Tr.

clinopyroxene Tr. Tr. 0

feldspar Tr. 18 16

amphibole 0 0 0

Fe–Ti oxide 0 0 0

(continued)


803

Dow



abundances of olivine and clinopyroxene correlate posi-tively with MgO (Fig. 3). The porphyritic basalts withmore than 10 vol. % phenocrysts contain coarse pheno-crysts, with sizes up to 15mm in the longest dimension(e.g. SH-15, 21, 25, 32, and 35 in Figs 3 and 4), whereasphenocryst-poor basalts (510 vol. % phenocrysts; e.g.SH-50 in Figs 3 and 4) generally have phenocrysts smallerthan 0·5mm.

Mineral compositionsWe present mineral compositions for the alkali basalts withMgO45wt %. These data are utilized to constrain theeffects of shallow magmatic processes and thus to estimatethe primary magma compositions.

Olivine

Olivines in St. Helena basalts are classified into two typesdepending on the absence or presence of kink banding(e.g. SH-25 in Fig. 4). The kink-banded olivine makes up0^57% of the total olivine phenocryst population in each

sample. The kink-banded olivine is mainly observed inhighly porphyritic samples, but is scarce in thephenocryst-poor basalts (510 vol. % phenocrysts).Figure 5 shows the olivine compositions for the samplesincluding the kink-banded olivines.Kink-banded olivines are commonly found in highly

porphyritic basalts, for example picrites and ankaramitesin oceanic islands, and are thought to be derived fromdisaggregated cumulates or mantle peridotite xenoliths(e.g. Albare' de & Tamagnan, 1988; Clague & Denlinger,1994; Garcia, 1996). In the St. Helena basalts, there areno systematic compositional differences between thekink-banded and kink-free olivines (Fig. 5). In addition,olivines in the porphyritic basalts have higher CaO con-tents and lower Mg# [100Mg/(MgþFe2þ)] comparedwith olivines in abyssal peridotites that have no signatureof metasomatism (Fig. 5). Therefore, kink-banded olivinesin the St. Helena basalts have a cognate origin with thekink-free olivines, and are not xenocrysts from mantleperidotite. The kink-banded olivines in the St. Helena bas-alts may be derived from highly crystalline portions of themagma conduits or crustal magma chambers, becauseshear strain caused by magma flow rotates minerals with-out plastic deformation only in a magma with low crystal-linity (e.g. Clarke et al., 2002).As shown in Fig. 6a, a range of olivine compositions

occurs even in single samples. Overall, olivine phenocrystshave core compositions of Mg#OL of 65^87 and showweak compositional zoning towards their rims (e.g.Fig. 5). Normally zoned olivines are most common,although reversely zoned crystals can also be found.Supplementary Data Appendix Fig. A1 shows the coreand rim compositions for the reversely zoned olivines inwhich the rims are a maximum of 5Mg# units higherthan the core.

Clinopyroxene

Clinopyroxene phenocrysts are euhedral, fractured, andembayed diopsides (Di44^49 En39^46 Fs8^15). They some-times form aggregates with olivine irrespective of the pres-ence or absence of kink band textures in the olivines.Clinopyroxene cores have Mg#Cpx of 74^90 (Fig. 6b)and titaniferrous outermost rims containing up to 5wt %TiO2. Clinopyroxenes exhibit more complicated compos-itional zoning than the olivines owing to the presence ofsector zoning and compositional heterogeneity aroundmulti-phase inclusions. Normally zoned olivines are mostcommon, but clinopyroxene phenocrysts showing weak re-verse zoning are also found (Supplementary DataAppendix Fig. A1).

Plagioclase

Plagioclase is minor or absent in St. Helena basalts withmore than 5wt % MgO (Fig. 2). Obvious exceptions to

Table 1: Continued

Reference standard compositionsBHVO-2 JB-2

GRM av. (n¼ 6) 1 SD GSJ av. (n¼ 3) 1 SD

ICP-MS (ppm)Sc 32 30·8 2·0 53·5 55·5 5·2Co 45 44·8 2·5 38 37·4 2·9Ni 119 119 7 16·6 13·9 0·8Cu 127 129 9 225 222 17Rb 9·11 9·02 0·36 7·37 6·33 0·31Sr 396 391 13 178 173 1Y 26 24·0 0·7 24·9 22·3 0·9Zr 172 174 5 51·2 47·0 1·3Nb 18·1 17·0 0·4 1·58 0·465 0·017Cs 0·1 0·095 0·004 0·85 0·789 0·006Ba 131 130 1 222 219 1La 15·2 15·1 0·2 2·35 2·23 0·02Ce 37·5 37·3 0·3 6·76 6·49 0·07Pr 5·35 5·26 0·04 1·01 1·12 0·01Nd 24·5 24·3 0·2 6·63 6·25 0·09Sm 6·07 6·06 0·08 2·31 2·24 0·04Eu 2·07 2·06 0·02 0·86 0·812 0·012Gd 6·24 6·27 0·09 3·28 3·21 0·06Tb 0·92 0·958 0·011 0·6 0·580 0·009Dy 5·31 5·32 0·05 3·73 3·94 0·05Ho 0·98 0·982 0·012 0·75 0·858 0·018Er 2·54 2·59 0·04 2·6 2·59 0·04Tm 0·33 0·332 0·005 0·41 0·377 0·009Yb 2 2·01 0·02 2·62 2·51 0·04Lu 0·274 0·277 0·004 0·4 0·384 0·008Hf 4·36 4·59 0·03 1·49 1·54 0·01Ta 1·14 1·12 0·01 0·13 0·031 0·001Tl 0·018 0·004 0·042 0·033 0·003Pb 1·6 1·58 0·28 5·36 5·01 0·07Th 1·22 1·21 0·02 0·35 0·252 0·002U 0·403 0·420 0·017 0·18 0·148 0·010

BTA, basaltic trachyandesite; Picro B, picrobasalt; TA, tra-chyandesite; TB, trachybasalt, Tr., trace (51 vol. %); n.d.,not determined. GRM and GSJ data are from Jochumet al. (2005) and Imai et al. (1995), respectively.


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this are three samples; SH-26, 32, and 39 (Fig. 2 andTable 1). In these samples, plagioclase is more abundantthan mafic minerals. The plagioclases have core compos-itions ranging from An62 to An84 and rim compositionsranging from An41 to An71 (Supplementary DataAppendix Table A3). Euhedral to subhedral plagioclasephenocrysts occasionally form aggregates with olivine andclinopyroxene phenocrysts.

Spinel

Olivine and clinopyroxene phenocrsyts enclose spinel in-clusions. The spinels show a wide range of Cr# [100Cr/(CrþAl)¼ 12^49] and Al2O3 contents (18^47wt %)(Supplementary Data Appendix Table A4), because thespinel inclusions appear both as discrete grains trapped ingrowing host minerals and as crystals precipitated fromtrapped melt. In the latter case, Al-rich and Cr#-poorspinel coexists with multiple phases such as Al- andTi-rich clinopyroxene, rho« nite, carbonate, and sulfide.

Groundmass

The groundmass consists of olivine, clinopyroxene, plagio-clase, alkali feldspar, Fe^Ti oxide, and glass. GroundmassFe^Ti oxides occasionally show spinel^ilmenite microinter-growths (lamellae), which indicate oxygen fugacities ran-ging from NNO (nickel^nickel oxide) to QFM ^ 2 (whereQFM is quartz^fayalite^magnetite) at temperatures of600^9008C (Supplementary Data Appendix Table A4)based on the equation of Carmichael (1967).

BULK-ROCK COMPOSIT IONSRock samples from lava flows, dikes, and clasts in the vol-caniclastic rocks range from picro-basalt through tra-chyandesite to trachyte and phonolite in composition(Fig. 7 and Table 1). There is no systematic compositionaldifference between the rocks from the older NE volcanoand the younger SW volcano (Fig. 7). Shield lava flowsand dikes exhibit a wide compositional range frompicro-basalt to trachyandesite. In contrast, late-stage

Fig. 2. Modal abundance of phenocrysts in St. Helena alkali basalts. Samples are arranged in order of decreasing MgO content from left toright. The actual MgO content is shown at the bottom of the bar graph. Phenocryst assemblages of porphyritic rocks gradually change fromwhere olivine and clinopyroxene are dominant to where olivine and plagioclase are dominant with decreasing MgO. Sample number isshown for representative samples.


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intrusions exposed on the SW volcano (Fig. 1) are mainlytrachytic and phonolitic in composition. On the NE vol-cano, trachyte is found only in clasts within the basal brec-cia sequence (e.g. SH-44A and 44B,Table 1).

St. Helena samples show a continuous compositionaltrend with a cusp at c. 5 wt % MgO for most of the majorelements (Fig. 8), suggesting that a change in the fractio-nating mineral assemblage controls the bulk-rock compos-itions. Ratios of elements with similar mineral^meltpartition coefficients in the major phenocryst minerals ofolivine, clinopyroxene, and plagioclase (e.g. Tb/Yb,Rb/Nb, and Zr/Sm) are nearly constant irrespective ofMgO content (Fig. 8). St. Helena basalts are characterizedby lower [Rb, Ba, Th, U, light rare earth element(LREE)]/Nb, and higher LREE/HREE (heavy REE)ratios than primitive mantle values (Sun & McDonough,1989) (Fig. 9).For comparison, Figs 8 and 9 also show the range of

HIMU OIB from three islands (Tubuai, Mangaia, andRurutu) in French Polynesia (Hanyu et al., 2011). Althoughthey include basalts with higher MgO contents (up to21wt %) and lower SiO2 (down to 43wt %) than the St.Helena basalts (which have maximum values of 16wt %MgO and 45wt % SiO2), both major and trace elementabundances overlap at equivalent MgO (10^16wt %)(Figs 8 and 9). This suggests that the St. Helena basaltsshare common geochemical characteristics with HIMUOIB from French Polynesia in terms of major and traceelements.

GROUNDMASS COMPOSIT IONSGroundmass compositions were calculated based on massbalance using the abundance and compositions of

Fig. 3. Modal abundance (vol. %) of mafic minerals vs bulk-rockMgO (wt %). Bulk-rock MgO contents show a positive correlationwith the abundance of clinopyroxene and olivine phenocrysts in therocks. Sample numbers indicate the samples shown in Fig. 4.

Fig. 4. Photomicrographs of representative St. Helena basaltic samples under plane-polarized light (SH-15, 21, 32, 35, and 50) and kinked olivinein a porphyritic basalt under cross-polarized light (SH-25). The sample numbers correspond to the labeled samples in Fig. 3. Three porphyriticbasalts (SH-15, 21, and 35) are characterized by coarse olivine and clinopyroxene, whereas SH-32 includes plagioclase in addition to coarse clino-pyroxene and olivine. Coarse olivine crystals in the porphyritic basalts often exhibit kink bands, as shown in SH-25. Phenocryst-poor basalt(SH-50) is characterized by rare olivine, plagioclase and clinopyroxene microphenocrysts.


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phenocrysts, together with bulk-rock compositions.Phenocrysts in the St. Helena basalts show largecompositional variations even in a single sample (Fig. 6),which easily propagates error on the estimation ofgroundmass compositions. To reduce the apparentcompositional variations of phenocrysts, we convertedthe phenocryst compositions from weight per cent tomole per cent. In addition, the two components FeO andMgO were treated as a single component (FeOþMgO).This treatment allows us to restrict the compositionsof olivine and clinopyroxene, and thus more accuratelycalculate the groundmass compositions on a mole percent basis.Open diamonds and open circles in Fig. 10b show the

groundmass compositions calculated from basalts having

bulk-rock MgO contents of 8^16wt % and 5^8wt %, re-spectively (Fig. 10a). The groundmass compositionsbroadly fall within the array defined by the bulk-rock com-positions. The groundmass compositions vary irrespectiveof bulk-rock compositions, because the influence of thephenocrysts on bulk-rock compositions differs betweensamples (Fig. 10c).

ORDER OF CRYSTALL I ZAT IONFor St. Helena basalts, the dominant phenocrysts are oliv-ine, clinopyroxene, plagioclase and spinel. Among these,the onset of clinopyroxene saturation can be assessed fromolivine chemistry. On a plot of CaO vs forsterite contentfor the cores of olivine phenocrysts, CaO first increases

Fig. 5. Compositions of kink-banded and kink-free olivine that occur in the same St. Helena basalt samples. Overlapping kink-banded andkink-free olivine compositions suggest that they are cognate. Olivines in the St. Helena basalts have higher CaO and lower Mg# than olivinesin abyssal peridotite.This indicates that the St. Helena basalts lack olivine xenocrysts derived from mantle peridotite. Data for abyssal peridotiteare from Arai & Fujii (1979), Sinton (1979), Fujii (1990), Juteau et al. (1990), Komor et al. (1990), Arai & Matsukage (1996), Niida (1997),Hellebrand et al. (2002), and Brunelli et al. (2006).


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with decreasing Mg#OL and then decreases beyondMg#OL of c. 85 (Fig. 11). The inflection in CaO asMg#OL varies may reflect a change in partition coeffi-cient and/or CaO activity in the melts. Because the CaOpartition coefficient shows a negative correlation with theMg#OL (Libourel, 1999), CaO in olivine should continu-ously increase with decreasing Mg#OL. This is inconsist-ent with what is observed, and thus a change in CaOactivity is more likely. The most probable phase to lowerCaO activity in the melt is clinopyroxene, which is thedominant phenocryst phase in the mafic St. Helena por-phyritic basalts (Fig. 2).Combining the inflection point at Mg#OL of 85

with the melt^olivine partition coefficient for Fe/Mg(KdFe/Mg) of 0·27^0·33, we conclude that clinopyroxenebegan to crystallize when the host melt reached anFeO/MgO of c. 1·1 (Mg# of c. 62), which correspondsto c. 11wt % MgO. This estimated saturation pointis also consistent with a weak inflection point in thevariation of NiO with the Mg#OL in the olivine cores(Fig. 12a).The fractionation of plagioclase may also decrease CaO

activity in the melt. However, plagioclase would begin tocrystallize when the melt reaches �5^8wt % MgO, wellafter clinopyroxene saturation. This is implied by the com-positional trend on an Al2O3 vs K2O plot (Fig. 10b).Figure 10b plots groundmass compositions (open symbols),together with bulk-rock compositions of sparsely phyric toaphyric rocks (55 vol. % phenocrysts, filled symbols). Theresultant trend is regarded as the liquid line of descent(LLD). The gentle slope of the LLD with greater than c.

1mol % K2O (continuous line in Fig. 10b) can be ex-plained by the involvement of plagioclase in the fractio-nated phases, because olivine and clinopyroxene containinsufficient Al2O3 to explain the LLD. In addition, whenwe focus on the trend in detail, the LLD with less thanc. 1mol % K2O seems to follow an olivine plus clinopyrox-ene control line (Fig. 10b). In considering this subtle differ-ence in the slope of the LLD, plagioclase would besaturated when the melt reaches 1mol % K2O. It should

Fig. 6. Mg# of olivine (a) and of clinopyroxene (b) vs FeO/MgO in bulk-rocks. Gray fields represent the mineral compositions in equilibriumwith the melt, assuming KdFe/Mg¼ 0·3�0·03 for olivine and 0·2�0·04 for clinopyroxene. Open circles represent mineral compositions for thesamples used in the fractionation correction (see text). The bulk-rock FeO/MgO is affected by the change in the amount of phenocrysts in themagma. For example, if the erupted magma has a lower crystal/melt ratio than the partially crystallized magma filling the magma chamberin a closed system, the bulk-rock FeO/MgO in the erupted magma increases.

Fig. 7. Total alkalis vs SiO2 diagram. Classification after Le Bas et al.(1986).Volcanic rocks on St. Helena show continuous variations alonga moderately alkaline trend. The bold dashed line represents theboundary between tholeiite and alkaline series rocks (Macdonald &Katsura, 1964).


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be noted that the K2O mol % corresponds to MgO of5^8wt %.The other phase that precipitated early from the basaltic

melts is spinel, which forms crystals that are enclosed inolivine and clinopyroxene phenocrysts. Because the olivine

around these spinel inclusions has a maximum Mg#OL

of 83, spinel probably became saturated before the growthof olivine with that composition. This remains true even ifwe consider the diffusion of Fe^Mg between spinel andolivine.

Fig. 8. Major oxides and element ratios vs MgO (wt %). The volcanic rocks from St. Helena show continuous trends with only small devi-ations. Inflections in the trends are observed in most oxides and trace elements at c. 5 wt % MgO. For comparison, HIMUOIB from three is-lands in French Polynesia (Tubuai, Mangaia, and Old-Rurutu, Hanyu et al., 2011) are also shown. The star symbols represent the primary meltcompositions of the St. Helena basalts corrected to Mg#OL of 87·6 and 90 (Table 2).


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DISCUSS IONEstimation of primary magmacompositionsVariations in bulk-rock compositions

The petrographical and geochemical variations of the St.Helena alkali basalts are mainly controlled by two factors.One is the compositional variation in the groundmass(Fig. 10b and c), which mainly reflects the extent of frac-tional crystallization of olivine, clinopyroxene, and plagio-clase. The other factor is the amount of phenocrysts.Bulk-rock MgO positively increases with the amount ofolivine and clinopyroxene (Fig. 3). Magma mixing has lessinfluence over the variations in the St. Helena basalts,although olivine and clinopyroxene occasionally showweak reverse zoning in several rock samples. A core of oliv-ine (Mg#OL of 75^83) is occasionally surrounded by arim with Mg# up to 5 units higher (Supplementary DataAppendix Fig. A1). This increase in Mg#OL can be ex-plained by a decrease in FeO/MgO in the melt by a max-imum of 0·5 units, considering a melt^olivine partitioncoefficient for Fe/Mg (KdFe/Mg) of 0·27^0·33. The expectedheterogeneity in melt compositions is much smaller thanthe variations in the St. Helena basalts with 5^16wt %MgO, which corresponds to a difference in FeO/MgO of2 units. Thus, differences in the extent of fractional crystal-lization and the amount of crystals entrained are largelyresponsible for the variations observed in the St. Helena

basalts. Ratios of elements with similar mineral^melt par-tition coefficients for the major phenocrysts, such asTb/Yb, Rb/Nb, and Zr/Sm (Fig. 8), are relatively constantin the St. Helena basalts, which suggests that these basaltswere derived from similar primary magmas.

Fractionation correction

We tested whether the bulk-rock compositions of porphy-ritic basalts represent liquid compositions. For three sam-ples with bulk-rock FeO/MgO of50·8 (SH-25, 35, and 86),the observed maximum Mg#OL of 83^87 in each sampleis less magnesian than the olivine that should crystal-lize from a liquid with this bulk-rock FeO/MgO (Fig. 6a).This is also the case for clinopyroxene (Fig. 6b). Thus,these three basalts, which include 23^51% mafic pheno-crysts (Table 1), have accumulated olivine andclinopyroxene.In contrast, the bulk-rock compositions of the five other

basalts (SH-24, 38, 45, 50, 84, and 90), which include6^39 vol. % phenocrysts, are in equilibrium with the mostMg-rich olivine observed in each sample in terms of Fe/Mg partitioning (open symbols in Fig. 6a). In addition,the calculated olivine compositions in equilibrium withtheir bulk-rock compositions plot within the field definedby the observed olivine core compositions on a Ni vsMg#OL plot (open symbols in Fig. 12b). Therefore, as afirst-order estimation their bulk-rock compositions can beregarded as original melt compositions. This suggests that

Fig. 9. Trace element concentrations normalized to primitive mantle values (Sun & McDonough,1989). Shaded area defines the compositionalrange of HIMU OIB from French Polynesia, which have compositions comparable with the St. Helena basalts at equivalent MgO contents(10^16wt %). Data for the HIMUOIB from French Polynesia are from Hanyu et al., 2011.


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the amount and compositions of the phenocrysts includedin these basalts would be appropriate to correct for the ef-fects of crystal differentiation that resulted in the evolvedgroundmass compositions along the liquid line of descent.The primary melts supplied to the St. Helena magmatic

system should be mafic enough to crystallize the mostMg-rich olivine phenocrysts (Mg#OL¼ 87·6) observedin the St. Helena basalts. Thus, we corrected for theeffect of crystal fractionation to estimate primary melt

compositions using the bulk-rock compositions of the fivebasaltic samples (SH-24, 38, 45, 50, 84, and 90).We did notuse the groundmass compositions for the fractionation cor-rection. This is because, for St. Helena basalts, the ground-mass compositions are relatively evolved and hard toaccurately calculate on a weight per cent basis by subtract-ing phenocrysts from the bulk-rock compositions owing tothe wide compositional variations in the phenocrysts,even in a single sample.

Fig. 10. Relationships between Al2O3 and K2O (both in mol %) for (a) bulk-rocks and (b) groundmass compositions. (b) also shows thebulk-rock compositions [same symbol as in (a)] of sparsely phyric to aphyric samples (55 vol. % phenocrysts). The compositional trend in(b) is regarded as the liquid line of descent of the melt. The regression line in (a) defined by the bulk-rock compositions of basalts with�8wt % MgO indicates olivineþ clinopyroxene control.The compositional trend of melts with55wt % MgO in (b) requires the involvementof plagioclase in the fractionated phase, in addition to olivine and clinopyroxene. (c) Relationship between olivine and clinopyroxene pheno-cryst abundance and the FeO*þMgO mol % of the bulk-rock and groundmass. The groundmass compositions were calculated based onmass balance using the phenocryst abundance, density of minerals (3·3 g cm�3 for olivine and pyroxene, 2·7 g cm�3 for plagioclase and ground-mass), and average phenocryst compositions. Also shown in (a) and (b) are the mineral core compositions for basaltic rocks with more thanc. 5 wt % MgO: olivine (Fo65�87), clinopyroxene (Di44^49 En39�46 Fs8�15) and plagioclase (An62�84).


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Because less plagioclase has fractionated from magmaswith MgO �8wt %, as discussed in the previous section,the fractionation correction was performed by incremen-tally adding (in steps of 0·01wt %) olivine and clinopyrox-ene in equilibrium with a given melt in the proportion of3:2 until the melt was in equilibrium with an olivine com-position of Mg#OL of 85. This ol:cpx mineral proportionwas used because it explains the trend in Ni vs Mg#OL

defined by olivines with Mg#OL 585, in rocks in whichboth olivine and clinopyroxene are saturated (Fig. 12a).Once the melt reaches the necessary composition to crys-tallize olivine with Mg#OL of 85, equilibrium olivinealone was added incrementally until the melt was in equi-librium with the most Mg-rich olivine (Mg#OL¼ 87·6)in St. Helena basalts. The equilibrium olivine compos-itions were calculated using the model proposed by

Fig. 11. CaO vs Mg# in the cores of olivine phenocrysts from two porphyritic basalts (SH-11 and 24). The inflection of CaO at Mg#OL of 85,indicated by the arrow, suggests the onset of clinopyroxene fractionation.

Fig. 12. (a) NiO vs Mg# for olivines in St. Helena basalts with more than 5wt % MgO.The observed olivine trend can be accounted for by achange in the fractionated phases. Thick continuous curves in (a) represent the fractionation of olivine, and olivine þ clinopyroxene. Thesewere calculated from the average composition of two samples SH-24 and 25. (b) Olivine compositions calculated to be in equilibrium with thebulk-rock compositions (diamond symbols). Open diamonds labeled with sample number are the samples used in the fractionation correction.Gray circles are the data from (a).


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Kinzler et al. (1990) and olivine stoichiometry. In contrast,the equilibrium clinopyroxene compositions were esti-mated using linear regression of the analyzed oxide con-tents (SiO2, TiO2, Al2O3, FeO, MnO, MgO, CaO, Na2O,K2O, and NiO) in the phenocrysts. This treatment allowsthe oxide contents of clinopyroxene to be expressed as afunction of its Mg# (Mg#Cpx). Then, equilibrium clino-pyroxene compositions were calculated for a given Mg#of melt, adopting the regression equations and partitioncoefficient (KdFe/Mg) value of 0·24 (Thy et al., 1991; Thy,1995).The effect of spinel fractionation was not considered in

this correction because the bulk-rock compositions of bas-alts with more than 5wt % MgO have constant Al2O3/TiO2 (Fig. 8) even though the spinels have differentAl2O3/TiO2 (Supplementary Data Appendix Table A4).Thus, spinel fractionation has a negligible effect on majoroxides and incompatible elements.Reconstructed primary melt compositions are listed in

Table 2 and shown in Fig. 8 (star symbols). The primarymelt has 14·6� 0·6wt % MgO (1 SD, standard deviation).The trace element compositions of the primary melt werecalculated based on the Rayleigh fractionation modelusing the mineral^melt partition coefficients listed inTable 2. Table 2 and Fig. 8 also show the primary meltcompositions that can coexist with olivine that hasMg#OL of 90, assuming the coexistence of the melt witha peridotite residue. This is because the highest Mg#OL

of 87·6 observed in the St. Helena basalts may only be aminimum value for the liquidus olivine crystallized fromthe primary melts. The estimated primary magmas inequilibrium with olivine that has Mg#OL of 90 containup to 20wt % MgO and their incompatible element con-centrations are c. 85% of those in compositions correctedto Mg#OL of 87·6.

Potential origin of the St. Helena primarymagmasFor the St. Helena primary magma, the minimum meltsegregation pressure can be constrained by lithosphericthickness (Ellam, 1992; Ito & Mahoney, 2005; Humphreys& Niu, 2009). The half-space lithosphere cooling model es-timates a lithospheric thickness of c. 60^70 km (Ito &Mahoney, 2005; Humphreys & Niu, 2009), combined withan age of 30 Ma for the seafloor around St. Helena at thetime of eruption of the lavas sampled in this study. Thepressure at the base of lithosphere may correspond to theminimum pressure at which the primary melt was segre-gated from its mantle source, if the overlying lithosphereprevents the ascent of a plume head.In considering a melt segregation pressure of more than

c. 2 GPa, the role of the garnet^clinopyroxene thermaldivide must be considered in the petrogenesis of St.Helena primary melt. This is stable at pressures above�2·7GPa in a system consisting of a simple mixture of

natural minerals (O’Hara, 1968) and above �2GPa fornatural rocks (Kogiso et al., 2004).To produce alkali basaltsthe source of the St. Helena primary melts should plot onthe silica-poor side of the garnet^clinopyroxene thermaldivide in (DiþJd)^Fo^Co (Yaxley & Green, 1998) orCaTs^Fo^Qtz (Kogiso et al., 2004) ternary diagrams. Inthis context, peridotite, Si-deficient pyroxenite and a peri-dotite^eclogite mixture with a high proportion of perido-tite are all potential sources of the St. Helena basalts.The compositions of the St. Helena primary melt are

compared with the products of melting experiments at2^7GPa using various starting materials in Fig. 13. Thisshows the compositional difference at equivalent MgOcontents (14^20wt %), relative to the St. Helena primarymelt calculated to be in equilibrium with olivine withMg#OL¼ 87·6^90. Two peridotitic starting materials(KR4003 and KLB-1), representative of pyrolite compos-itions (McDonough & Sun, 1995), yield partial melts withhigher SiO2 and Al2O3 (Fig. 13a), and lower FeO*(Fig. 13b) than the St. Helena primary melt. CO2 is prob-ably involved in the genesis of the St. Helena basalts be-cause carbonate phases are observed in trapped meltinclusions in olivine phenocrysts. However, partial meltsof CO2-bearing peridotite (Dasgupta et al., 2007) cannotreproduce the CaO contents of the St. Helena primarymelt (Fig. 13c). Simple melting of primitive mantle orMORB-source mantle is also ruled out in terms of traceelements and isotopic ratios. Partial melts of primitivemantle (Rb/Nb40·89, Sun & McDonough, 1989) cannotexplain the low Rb/Nb (�0·4) of the St. Helena primarymelt, as long as common mantle minerals participate inthe melting. In addition, depleted MORB-source mantle(DMM) alone is an unlikely source, because the St.Helena basalts have higher 206Pb/ 204Pb (420·5, Chaffeyet al., 1989) and 187Os/188Os (up to 0·14^0·20, Reisberget al., 1993), and lower 3He/4He (5^7 RA, Hilton et al.,2000) than modern MORB.Partial melts of garnet pyroxenite (MIX1G, e.g. Kogiso

et al., 2003;77SL-582, Keshav et al., 2004), olivine websterite(HK66, Hirose & Kushiro, 1993) and a mixture of basalt^fertile peridotite in the proportions of 3:7 (KG2) and 1:1(KG1) (Kogiso et al., 1998) also fail to reproduce the SiO2,FeO*, CaO, Al2O3, Na2O and TiO2 contents of the St.Helena primary melt (Fig. 13). Reaction between a partialmelt of hornblendite and peridotite at 1·5GPa can broadlyexplain the major element compositions of some alkalicOIB with 44^47wt % SiO2 (Pilet et al., 2008). However,this is not the case for the St. Helena basalts. The experi-mental melts contain higher Al2O3, by c. 2^4wt %, andlower MgO, by 2^5wt %, and FeO, by 3^4wt %, thanthe St. Helena primary melt at an equivalent SiO2 contenton an anhydrous basis (not shown).In contrast, two experimental studies have reported

melt compositions broadly similar to the St. Helena


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Table 2: Calculated primary melt compositions and partition coefficients used in estimating the primary melt compositions

Primary melt compositions (Mg#OL¼ 87·6) Mg#OL¼ 90

SH24 SH38 SH45 SH50 SH84 SH90 average 1 SD average 1 SD

wt %

SiO2 44·83 45·26 44·90 45·37 45·04 44·41 44·97 0·34 44·32 0·29

TiO2 2·55 2·19 2·52 2·61 2·49 2·32 2·45 0·16 2·08 0·14

Al2O3 10·52 12·63 11·93 12·05 11·74 11·28 11·69 0·72 9·95 0·63

FeO* 12·25 12·01 11·64 12·04 12·25 12·70 12·15 0·35 11·94 0·29

MnO 0·19 0·18 0·17 0·19 0·19 0·22 0·19 0·02 0·18 0·02

MgO 14·76 14·32 13·68 14·35 14·65 15·57 14·56 0·62 19·60 0·71

CaO 11·51 10·40 11·82 10·12 10·45 11·00 10·88 0·68 9·27 0·62

Na2O 2·21 2·04 2·19 1·95 2·05 1·61 2·01 0·22 1·71 0·19

K2O 0·80 0·63 0·77 0·85 0·76 0·56 0·73 0·11 0·62 0·09

P2O5 0·39 0·33 0·38 0·47 0·39 0·31 0·38 0·06 0·32 0·05

Total 100 100 100 100 100 100

ppm

Rb 14·7 11·6 14·6 14·3 14·2 10·0 13·2 2 11·1 2

Ba 205 187 200 242 200 169 200 22 169 29

Th 2·67 2·46 3·06 3·58 2·93 2·39 2·85 0·4 2·40 0·5

U 0·798 0·693 0·875 0·981 0·830 0·652 0·805 0·1 0·679 0·1

Nb 32·6 29·5 34·8 42·1 35·6 26·9 33·6 5 28·4 6

Ta 2·09 1·87 2·26 2·72 2·23 1·75 2·15 0·3 1·82 0·4

La 25·2 23·9 26·5 32·9 26·9 21·37 26·1 4 22·1 5

Ce 53·7 47·9 55·9 69·7 56·6 45·3 54·9 8 46·6 10

Pb 1·30 1·50 1·27 1·72 1·29 1·0 1·35 0·2 1·14 0·3

Pr 6·63 6·00 6·97 8·62 6·90 5·60 6·79 1·0 5·78 1·3

Sr 418 433 418 464 427 378 423 25 356 41

Nd 27·7 25·0 28·8 36·1 28·1 23·56 28·2 4 24·1 6

Sm 5·90 5·28 6·19 7·81 5·90 5·10 6·03 0·9 5·19 1·3

Zr 172 149 178 215 174 140 171 24 146 33

Hf 4·40 3·84 4·66 5·67 4·34 3·77 4·45 0·6 3·82 0·9

Eu 1·90 1·77 1·99 2·41 1·91 1·66 1·94 0·2 1·66 0·4

Gd 5·61 5·29 5·84 7·20 5·56 4·87 5·73 0·7 4·93 1

Tb 0·831 0·774 0·882 1·101 0·840 0·732 0·860 0·1 0·743 0·2

Dy 4·36 4·11 4·68 5·72 4·41 3·89 4·53 0·6 3·91 1·0

Y 18·4 19·8 20·3 24·3 19·5 16·6 19·8 2 17·1 4

Ho 0·776 0·779 0·857 1·045 0·806 0·705 0·828 0·1 0·716 0·2

Er 2·02 2·06 2·26 2·72 2·15 1·86 2·18 0·3 1·9 0·4

Yb 1·49 1·55 1·77 2·03 1·67 1·42 1·66 0·2 1·43 0·3

Lu 0·211 0·221 0·250 0·294 0·243 0·206 0·237 0·03 0·205 0·05

Ni 465 445 350 362 358 429 401 46 885 101

87Sr/86Sr 0·70289 0·00013

143Nd/144Nd 0·51287 0·00009

206Pb/204Pb 20·7 0·3

207Pb/204Pb 15·77 0·05

208Pb/204Pb 40·0 0·2

176Hf/177Hf 0·28288 0·00009

Added phase (wt %)

Cpx 0 0 0 0·096 0 0

Ol 0·09 0·11 0·08 0·17 0·09 0·12

(continued)


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primary melt. One is a partial melt of an Fe-rich garnetlherzolite (PHN 1611; Fig. 13) at 1470^14808C and 3GPa(Kushiro, 1996). It should be noted that PHN 1611, asheared lherzolite xenolith from kimberlite at ThabaPutsoa, has been affected by melt infiltration and mechan-ical mixing of inhomogeneous peridotite (Smith et al.,1993). The other is a melt derived from the reaction be-tween a partial melt of MORB-like coesite eclogite (GA1)and pyrolitic peridotite (MPY90) at 1425^15008C and3GPa (Yaxley & Green, 1998; Fig. 13). At equivalent MgOcontents, several melt compositions reported from thetwo studies (Kushiro, 1996; Yaxley & Green, 1998) fallwithin 10% of the average St. Helena primary melt forSiO2, Al2O3, CaO, and FeO* (Fig. 13). The broad compos-itional similarity between the St. Helena primary meltand the experimentally obtained melts suggests that anadditional component other than peridotite plays an im-portant role in producing the St. Helena primary melt.Such a component can be introduced into the source ofthe St. Helena basalts either through earlier melt-relatedmetasomatism or by instantaneous melt reaction withperidotite.

Evaluation of the oceanic crust-recyclinghypothesisAncient subducted oceanic crust is one of the most likelycomponents that have contributed to the genesis of OIB(Hofmann & White, 1982; Weaver, 1991; Chauvel et al.,1992; Stracke et al., 2003, 2005; Sobolev et al., 2005;Willbold & Stracke, 2006; Prytulak & Elliott, 2007).Fractionation between trace elements in the oceanic crustduring seafloor alteration (e.g. Chauvel et al., 1992),subduction-related processes (Zindler & Hart, 1986;Weaver, 1991), and partial melting in the lower mantle(e.g. core^mantle boundary, Kogiso et al., 1997; Tauraet al., 2001; Hanyu et al., 2011) have all been proposed as po-tential processes to explain the geochemical characteristicsof HIMUOIB. However, further data are needed to valid-ate the details of the oceanic crust-recycling hypothesis.This is because the conclusions of previous studies arestrongly dependent on assumptions about element mobilityduring the alteration and dehydration of the subducted ig-neous oceanic crust (Weaver, 1991; Chauvel et al., 1992;Kogiso et al., 1997). In addition, for simplicity, several stu-dies neglected to consider the influence of peridotite melt-ing on the isotopic ratios and incompatible element ratiosof HIMU OIB (e.g. Weaver, 1991; Hauri & Hart, 1993;Stracke et al., 2003).Unlike most of the previous studies, we have made an in-

verse calculation of the recycled oceanic crust composition,together with the petrogenetic conditions (e.g. meltingdegree, melting mode, amount of melt contribution fromthe recycled oceanic crust) for the St. Helena basaltsource. This calculation does not need to assume thefluid-mobile element abundances in the recycled oceanic

Table 2: Continued

Partition coefficients Peridotite

OL Cpx PM DMM

wt %

SiO2 44·90 44·71

TiO2 0·20 0·13

Al2O3 4·44 3·98

FeO* 8·03 8·18

MnO 0·13 0·13

MgO 37·71 38·73

CaO 3·54 3·17

Na2O 0·36 0·28

K2O 0·03 0·01

P2O5 0·00 0·00

Total 99·34 99·32

ppm

Rb 0·0003 0·011 0·635 0·05

Ba 0·000005 0·022 6·989 0·563

Th 0·000007 0·022 0·085 0·0079

U 0·000009 0·016 0·021 0·0032

Nb 0·00005 0·026 0·713 0·1485

Ta 0·00005 0·063

La 0·0002 0·18 0·687 0·192

Ce 0·00007 0·31 1·775 0·55

Pb 0·0003 0·06 0·18 0·018

Pr 0·0003 0·46 0·276 0·107

Sr 0·00004 0·104 21·1 7·664

Nd 0·0003 0·63 1·354 0·581

Sm 0·0009 0·96 0·444 0·239

Zr 0·001 0·52 11·2 5·082

Hf 0·0029 0·9 0·309 0·157

Eu 0·0005 0·86 0·168 0·096

Gd 0·0011 1·02 0·596 0·358

Tb 0·0019 1·18

Dy 0·0027 1·16 0·737 0·505

Y 0·0082 1·14 4·55 3·328

Ho 0·01 1·19 0·164 0·115

Er 0·0109 1·1 0·48 0·348

Yb 0·024 0·9 0·493 0·365

Lu 0·02 0·97 0·074 0·058

Ni87Sr/86Sr 0·70500 0·70230143Nd/144Nd 0·512638 0·513250206Pb/204Pb 17·72 17·72207Pb/204Pb 15·49 15·49208Pb/204Pb 37·36 37·36176Hf/177Hf 0·28280 0·28326

*Total Fe given as FeO.Partition coefficients for olivine are from Halliday et al.(1995). Partition coefficients for clinopyroxenes are fromphenocryst/matrix partition coefficient reported byD’Orazio et al. (1998) for hawaiite samples. Partition coef-ficients of Tb for olivine and Gd for clinopyroxene are set tomean values for Gd and Dy, and Eu and Tb, respectively.Isotopic ratios of St. Helena basalts from compilation byStracke et al. (2003). Primitive mantle (PM) and depletedMORB mantle (DMM) are from Sun & McDonough (1989)and Workman & Hart (2005), respectively.


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crust, which is the advantage of this type of modeling.By comparing the calculated recycled crust with observedfresh MORB compositions, we can evaluate whetherchemical modification involving seafloor alteration and

slab dehydration played a primary role in producing thedistinctive HIMU geochemical signature of the St.Helena basalts. Below, we provide an outline of our geo-chemical modeling.

Fig. 13. Compositional difference between the St. Helena primary melt and experimental melts at an equivalent MgO content (15^20wt %).Experimental melts are normalized on an anhydrous basis. To calculate the compositional difference, each experimental melt composition hadsubtracted from it the St. Helena primary melt composition, corrected using Mg#OL of 87·6^90. Melt compositions from experiments at2^7GPa are from Hirose & Kushiro (1993), Kushiro (1996), Kogiso et al. (1998, 2003),Walter (1998), Yaxley & Green (1998), Hirschmann et al.(2003), Keshav et al. (2004), and Dasgupta et al. (2007).


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Modeling set-up

The basic set-up of our model is shown schematicallyin Fig. 14. First, oceanic crust is recycled back intothe mantle at a subduction zone at some time in thepast. The isotopic ratios in the subducted oceanic crustevolve depending on its residence time in mantle andthe P/D ratios. We assume that the subducted oceaniccrust is entrained in an upwelling mantle plume. In thefollowing discussion, the term ‘recycled crust’ is used forthe subducted igneous oceanic crust entrained in themantle plume. The recycled crust is represented bya MORB composition to calculate reliably the com-position of the recycled crust-derived melt usingexperimentally constrained melting relationships(Pertermann & Hirschmann, 2003a). The solidus tem-perature is positively and negatively correlated with theMg# and alkali content of the source material, respect-ively (Walter & Presnall, 1994; Herzberg et al., 2000;Hirschmann, 2000; Kogiso et al., 2004). Thus, duringmantle upwelling, the MORB-like recycled crust wouldbegin to melt before the onset of host peridotite melting(Yasuda et al., 1994; Yaxley & Green, 1998; Kogiso et al.,2004).The fate of the recycled crust-derived melt is a critical

issue, which affects how OIB melts are generated. Becausethe recycled crust-derived melt is in disequilibrium withthe host peridotite, this Si-saturated melt will react withthe peridotite to produce various lithologies, ranging fromperidotite to pyroxenite, and the compositions of the re-maining melt will change as indicated by reaction experi-ments (Sekine & Wyllie, 1983; Sen & Dunn, 1995; Yaxley& Green, 1998).We simplify the interaction between the recycled

crust-derived melt and surrounding peridotite in anupwelling mantle plume to two separate stages (Fig. 14a).First, the recycled crust-derived melt is mixed withthe peridotite to form a hybrid source. Then, the hybridsource is partially melted to generate the St. Helenaprimary basalt during further upwelling of the mantleplume. This simplification is required because thechanges in both phase assemblage and phase proportionsduring continuous reaction in an upwelling heterogeneousmantle plume are difficult to constrain. A morelimited melt^solid interaction model, rather than our com-plete mixing model, is another end-member option, ifmelt^solid reaction is inhibited (e.g. by formation of opxin the wall-rocks), or if the melt velocity is high (e.g.Stracke & Bourdon, 2009).We do not investigate this morecomplex process here. If it is applicable, however, the com-positions of OIB may simply be approximated by a mix-ture of two kinds of melt derived from recycled crust andperidotite (Ito & Mahoney, 2005; Stracke & Bourdon,2009).

Modeling parameters

The mixing and melting events within the upwellingmantle plume are described using the free variables listedin Table 3 and illustrated in Fig. 14a. Melting of therecycled crust is expressed by an accumulated fractionalmelting model [equation (A1) in the Appendix]. Theweight fraction of melt produced is represented by FRC,which is allowed to vary between 0·2 and unity (see theAppendix).Within the range of FRC the residue consists ofclinopyroxene and/or garnet (Pertermann & Hirschmann,2003a). Proportions of the residual clinopyroxene andgarnet are parameterized as a function of FRC [equations(A2) and (A3) in the Appendix] based on the phase pro-portions reported in the melting experiments forMORB-like basalt (Pertermann & Hirschmann, 2003b).Sub-solidus phase proportions of recycled crust weretaken from Pertermann & Hirschmann (2003b) to calcu-late the bulk partition coefficients before melting(FRC¼ 0). The element abundances in the recycledcrust-derived melt are thus controlled by FRC and the min-eral^melt partition coefficients listed inTable 4.The composition of the hybrid source is assumed to be a

mixture of the recycled crust-derived melt and the modelperidotite [equation (A4) in the Appendix]. The mixingratio is expressed as a weight fraction of the model perido-tite component in the hybrid source mantle (XPeri). XPeri

is limited to values of 0·5^1, because this is required to pro-duce a Si-deficient source through the mixing of peridotiteand the Si-saturated melt derived from a MORB-likebasalt on the garnet^clinopyroxene thermal divide.The model peridotite composition is varied to test for

the effects of possible peridotite heterogeneity, and isapproximated by a mixture of primitive mantle (PM, Sun& McDonough, 1989) and depleted MORB source mantle(DMM, Workman & Hart, 2005) [equation (A5) in theAppendix]. The mixing proportion between PM andDMM is controlled by a free variable of XDMM, which rep-resents the weight fraction of the DMM component in themodel peridotite (Table 3). Trace element abundances andpresent-day isotopic ratios of the PM and DMM are listedinTable 2.Melting of the hybrid source is expressed using a batch

melting model [equations (A6) and (A7) in theAppendix]. In this model, the weight fraction of the meltproduced by melting of the hybrid source (FOIB) is a vari-able that depends on the free parameters XDMM, XPeri

and FRC, as well as the bulk partition coefficient of Nb[equations (A1)^(A6) in the Appendix]. Bulk partition co-efficients for the hybrid source melting are calculatedusing the partition coefficients listed in Table 4 and themineral abundance in the residual hybrid source. The re-sidual mineral abundances can also be treated as unknownvariables. There are four potential residual phases; olivine(OL), orthopyroxene (Opx), clinopyroxene (Cpx), and


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garnet (Grt). The weight fractions of three mineral phasesin the residue (XOpx, XCpx, and XGrt) are treated fully asfree variables, but that of olivine (XOL) is expressed as aparameter that depends on XPeri, FRC, FOIB, XOpx, XCpx,and XGrt based on the mass balance for SiO2þ0·5Al2O3

þ 0·5Na2O mol % [equation (A13) in the Appendix].

Overall, the trace element abundances in the model OIBare varied by the free parameters XDMM, XPeri, FRC, andmineral abundances in the residual hybrid source (XOpx,XCpx, and XGrt) (Fig. 14 and Appendix Fig. A1).Based on the P/D ratios in the recycled crust, our model

can also test the isotopic evolution of subducted

Fig. 14. Schematic illustration of the geochemical model that postulates that subducted MORB-like oceanic crust is entrained by a mantleplume. The magmatic processes that occur in the upwelling mantle plume are simplified as follows: (1) the recycled oceanic crust is partiallyor totally melted; (2) the recycled crust-derived Si-saturated melt is mixed with the host peridotite to form a Si-deficient hybrid source; (3) par-tial melting of the hybrid source results in the St. Helena primary melt. In (a), mass-balance relations are shown using XPeri, FRC, and FOIB.(b) schematically shows the change in 143Nd/144Nd vs Nd abundance during the modeling. The initial isotopic composition of the recycledcrust is estimated using the two free parameters, DMM age and recycling age (c; see text). The recycling age represents the residence time ofthe subducted oceanic crust in the mantle. Modeling consists of four calculation steps (see text for details). At the third calculation step,non-conservative element abundances in the recycled crust are calculated to satisfy the element abundance of the St. Helena primary melt[gray field in (d)]. In the fourth calculation step, the recycled crust compositions are further constrained to satisfy the isotopic compositions ofSt. Helena basalts [black field in (d)].


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oceanic crust. Initial isotopic ratios of the recycled crustare estimated using two free parameters; the DMM ageand the recycling age, which represent the age whenDMM and the oceanic crust were separated from BulkSilicate Earth (¼ PM) and DMM, respectively (Fig. 14c).The method to calculate the initial isotopic ratios basicallyfollows that proposed by Stracke et al. (2003) (see theAppendix). Because the recycling age represents the resi-dence time of the subducted oceanic crust in the mantle(Fig. 14c), the present-day isotopic ratios of the recycledcrust vary depending on its P/D ratio, the DMM age, andthe recycling age. Mixing of the recycled crust-derivedmelt with the model peridotite determines the isotopicratios of the model OIB [equation (A14) in the Appendixand Fig. 14b]. As a result, the isotopic ratios of the modelOIB are affected by (1) the present-day isotopic ratios ofthe model peridotite controlled by XDMM, (2) thepresent-day isotopic ratios of the recycled crust controlledby its P/D ratios, DMM age, and recycling age, (3) XPeri,and (4) the element abundances in the peridotite andrecycled crust-derived melt (e.g. Sr, Nd, Hf, and Pb).Details for the solution of this model are given below.

Calculation procedures

Eight free variables (XDMM, XPeri, FRC, XOpx, XCpx, XGrt,DMM age, and recycling age) and the abundance of 10non-conservative elements (Rb, Ba, Th, U, Pb, La, Ce, Pr,Nd, and Sm) in the recycled crust are unknown in thismodel. These unknown values are calculated through thefour calculation steps (see Appendix, Fig. A1). In the firstand second steps, the combination of six free parameters(XDMM, XPeri, FRC, XOpx, XCpx, and XGrt) is optimized toreproduce the abundances of 11 conservative incompatibletrace elements (Nb, Zr, Hf, Y, Eu, Gd, Dy, Ho, Er, Yb,

and Lu) in the St. Helena primary melt and to satisfy thephase equilibria. The third and fourth steps constrain thetwo free variables of DMM age and recycling age, and cal-culate the abundance of 10 non-conservative elements inthe recycled crust, to reproduce the element abundancesand isotopic ratios of the St. Helena primary melt.

First step: calculation using conservative elements

In the first step, we use only the conservative elementabundances. The conservative elements are assumed to beless mobile during seafloor alteration and dehydration ofthe oceanic crust. During alteration, Rb, U, and occasion-ally Th and Ba are mobile, but the conservative elementsare immobile (Staudigel et al., 1996; Alt & Teagle, 2003;Bach et al., 2003; Kelley et al., 2003). On the other hand,during subduction, this assumption may only be validwhen the slab surface geotherm is moderate to relativelycool. This is an important limitation of this modeling.According to the experiments of Kessel et al. (2005) andKogiso et al. (1997), this assumption is fulfilled (partitioncoefficient, Dmineral/fluid �1), when the slab geotherm islower than 800^9008C at 4^6GPa. Our model, thus, ex-cludes the modification of the oceanic crust by slab meltingbeneath arcs.Because the conservative elements are assumed to be less

mobile during seafloor alteration and slab dehydration,their concentration in the recycled crust can be approxi-mated by that of fresh oceanic crust. Here, the conserva-tive elements in the recycled crust are set using the valuesfor fresh N-MORB from Hofmann (1988). Once therecycled crust compositions are given, the element abun-dance in the model OIB melt is calculated using six freeparameters (XDMM, XPeri, FRC, XOpx, XCpx, and XGrt)through the melting and mixing processes (Fig. 14).

Table 3: Modeling parameters

Symbol Parameter Range of value

FOIB MOIB/MSource dependent

FRC MRCmelt/MRC 0·2–1

XPeri MPeri/(MPeriþMRCmelt) 0·5–1

XOL Weight fraction of olivine in the residual hybrid source dependent

XOpx Weight fraction of orthopyroxene in the residual hybrid source 0–1

XCpx Weight fraction of clinopyroxene in the residual hybrid source 0–1

XGrt Weight fraction of garnet in the residual hybrid source 0–1

XDMM MDMM/(MDMMþMPM) 0–1

DMM age The age when DMM was separated from Bulk Silicate Earth 2–3·4 Ga

Recycling age The age when oceanic crust was produced from DMM 40·1 Ga

FOIB is dependent on XDMM, FRC, XPeri, and bulk partition coefficient of Nb. XOL is dependent on XPeri, FRC, FOIB, XOpx,XCpx, and XGrt. MPeri, MDMM, MPM, MRC, MRCmelt, MSource, and MOIB are the mass of model peridotite, DMM, PM,recycled crust, recycled crust-derived melt, hybrid source, and model OIB, respectively. It should be noted that therecycling age should be less than the DMM age (see text).


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In practice, we varied three parameters to cover the pos-sible range (XDMM 0^1, XPeri 0·5^1, and FRC 0·2^1), andthen optimized the residual mineral abundance (XOpx,XCpx, and XGrt; XOL is constrained by the other param-eters, as described above) by minimizing the differencesbetween the abundances of the conservative elements inthe model OIB and the St. Helena primary melts(Appendix Fig. A1). The combinations of six free param-eters are accepted when the model OIB compositions fallwithin 2 SD of the mean St. Helena primary melt(Table 2).We use the St. Helena primary melt compositionin equilibrium with Mg#OL of 87·6 (Table 2). However,differences in the primary melt compositions do not affectour main conclusions.

Second step: constraint of modeling parameters usingmajor elements

The mineral abundances in the residual hybrid source arestrictly related to the major element compositions of the

source mantle and the P^T conditions of melting in thenatural system. Consequently, the mineral abundancesoptimized in the first step (Fig. 15a) are further con-strained based on phase equilibria. For this purpose, themineral abundances in the residual hybrid source are cal-culated using the thermodynamic model pMELTS(Ghiorso et al., 2002) for the various combinations of par-ameters obtained from the first calculation step.Major element abundances in the hybrid source are

calculated using the compositions of the peridotiteand partial melts of MORB-like eclogite, as well asXDMM and XPeri based on equation (A4) in the Appendix.The compositions of PM and DMM are fromMcDonough & Sun (1995) and Workman & Hart (2005),respectively (Table 2). In contrast, the composition of thepartial melt from MORB-like eclogite was interpolatedusing the experimental data at 2^3GPa (Pertermann &Hirschmann, 2003a) for the melt fraction equivalentto FRC.

Table 4: Partition coefficients used for melting of recycled crust and hybrid source

Melting of recycled crust Melting of hybrid source

Cpx Grt Rutile Coesite Ol Opx Cpx Grt

Rb 0·0035 0·0007 0 0 0·0003 0·0002 0·0004 0·0002

Ba 0·0019 0·00004 0 0 0·000005 0·000006 0·0004 0·00007

Th 0·007 0·0015 0 0 0·00005 0·002 0·0059 0·009

U 0·008 0·006 0 0 0·00038 0·002 0·012 0·028

Nb 0·021 0·008 22 0 0·0005 0·004 0·015 0·015

La 0·0339 0·0017 0 0 0·0005 0·004 0·03 0·0007

Ce 0·064 0·0061 0 0 0·0005 0·004 0·08 0·017

Pb 0·13 0·18 0 0 0·003 0·009 0·012 0·005

Pr 0·11 0·0196 0 0 0·00046 0·008 0·15 0·0405

Sr 0·08 0·005 0 0 0·00004 0·0007 0·091 0·0007

Nd 0·1627 0·0564 0 0 0·00042 0·012 0·2 0·064

Sm 0·3 0·2741 0 0 0·0011 0·02 0·299 0·23

Zr 0·093 0·4 4·1 0 0·0011 0·02 0·2835 0·4

Hf 0·17 0·31 6 0 0·0011 0·024 0·2835 0·4

Eu 0·3574 0·4883 0 0 0·0011 0·0425 0·3245 0·715

Ti 0·36 0·21 9 0 0·015 0·086 0·35 0·6

Gd 0·44 0·817 0 0 0·0011 0·065 0·35 1·2

Dy 0·5323 1·9035 0 0 0·0027 0·065 0·4 2

Y 0·58 2·9 0 0 0·00785 0·065 0·41 2·5

Ho 0·5683 2·5961 0 0 0·010425 0·065 0·415 2·75

Er 0·59 3·3047 0 0 0·013 0·065 0·42 3

Yb 0·5987 4·5608 0 0 0·02 0·08 0·45 5

Lu 0·5915 5·0435 0 0 0·02 0·12 0·45 6

Ol, olivine; Opx, orthopyroxene; Cpx, clinopyroxene; Grt, garnet. Data for recycled crust melting are from Klemme et al.(2002) for clinopyroxene and garnet, and Klemme et al. (2005) for rutile. Data for hybrid source melting are from Johnson(1998) Salters et al. (2002), and Salters & Stracke (2004).


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By using the hybrid source mantle compositions, the re-sidual mineral abundances are calculated using pMELTSfor various FOIB, assuming melting pressures of 2^3GPa.The melting pressure is chosen to maintain internal con-sistency with the experimental conditions for MORB-likeeclogite (GA1, Pertermann & Hirschmann, 2003a). Itshould be noted that this also agrees with the pressure atwhich the St. Helena primary melt was probably segre-gated from the mantle residues as described above. The re-sidual mineral abundances calculated using majorelements can be directly compared with the results fromthe previous step for the common combinations of XDMM,XPeri, FRC, and FOIB (Fig. 15).To investigate the accuracy of pMELTS, we simulated

the partial melting of a depleted peridotite (DMM1,Wasylenki et al., 2003) at 1GPa and compared the resultswith the near-solidus melting experiment (melting degree41·6%,Wasylenki et al., 2003). At the same degree of melt-ing, the mineral abundances from pMELTS deviate fromthe experimental results within 20% for the abundant re-sidual minerals (10^70wt %), and up to 60% for theminor minerals (510wt % in residue). By considering theuncertainties between the experimental and pMELTS re-sults, we allow 20% and 50% errors for the mineral abun-dances of moderately abundant minerals (XOL, XOpx, andXCpx) and the minor mineral (XGrt), respectively. Theseerror windows were used to filter the calculation resultsfor XOL, XOpx, XCpx, and XGrt from the conservative elem-ent mass-balance calculations performed in the first step(Fig. 15a). The filtered mineral abundances are shown byopen symbols in Fig. 15b. Although the first step calcula-tions give a wide range of XPeri (0·5^1·0), constraints fromthe pMELTS mineralogy (shaded areas in Fig. 15b)reduce the range in XPeri to 0·82^0·95.Based on the above, the accepted modeling parameters

are summarized as follows: XDMM of 0^1, XPeri of0·82^0·95, FRC of 0·7^1·0, XOL of 0·46^0·56, XOpx of0·16^0·23, XCpx of 0·17^0·28, and XGrt of 0·06^0·10. Thesecombinations of parameters are used in the next calcula-tion steps. It should be noted that some parameters (e.g.XPeri40·92 and XDMM50·7) are rejected in the third andfourth calculation steps because they require negativeelement abundances in the recycled crust or fail to repro-duce the observed isotopic ratios of St. Helena basalts.

Third step: calculation using non-conservative elements

By using the St. Helena primary melt composition(Table 2) and the optimized combination of six free param-eters, non-conservative element abundances in the recycledcrust are calculated following the reverse path of themodel (melting of hybrid source, mixing of model perido-tite and recycled crust-derived melt, and melting ofrecycled crust, in that order). The recycled crust compos-ition calculated using equations (A1)^(A7) in theAppendix exhibits a certain compositional range reflecting

the variability inherited from the composition ofSt. Helena primary melt (average �1 SD), which isshown schematically by the gray shaded square field inFig. 14d.

Fourth step: filtering by isotopic ratios

The calculated non-conservative element abundances inthe recycled crust and the combination of six free param-eters are further filtered to satisfy the isotopic ratiosobserved in the St. Helena basalt (average �1 SD,Table 2). In this step, the following calculations are repeat-edly performed for each combination of six free param-eters (XDMM, XPeri, FRC, XOpx, XCpx, and XGrt) and therecycled crust compositions obtained in the third step.First, the initial Sr, Nd, Pb, and Hf isotopic ratios of the

recycled crust are calculated using the two free parametersof the DMM age and recycling age. Then, the present-dayisotopic ratios of the recycled crust are calculated usingthe recycling age and P/D ratios of the recycled crust(Rb/Sr, Sm/Nd, U/Pb, and Th/Pb) obtained in the thirdstep (Fig. 14b). It should be noted that the Lu/Hf ratio ofthe recycled crust is same as that of average MORB(Hofmann, 1988), because Lu and Hf are assumed to beconservative elements. Next, the 87Sr/ 86Sr, 143Nd/144Nd,206Pb/ 204Pb, 208Pb/ 204Pb, and 176Hf/177Hf isotopic ratiosof the model OIB, which are equivalent to that of thehybrid source, are calculated by mixing the model perido-tite with the recycled crust-derived melt using XDMM,XPeri, and FRC [equation (A14) in the Appendix]. This cal-culation was carried out to cover the possible range ofDMM age (2^3·4 Ga) and recycling age (0·1^3·4 Ga). Itshould be noted that the recycling age should not exceedthe DMM age (see Appendix). The isotopic ratios of themodel peridotite are calculated using XDMM and the iso-topic ratios of PM and DMM (Table 2).Based on these calculations, we accepted the combin-

ation of all independent free parameters (XDMM, XPeri,FRC, XOpx, XCpx, XGrt, DMM age, and recycling age) andthe non-conservative element abundances in the recycledcrust (Table 5), when the model OIB reproduced thetrace element abundances and 87Sr/ 86Sr, 143Nd/144Nd,206Pb/ 204Pb, 208Pb/ 204Pb, and 176Hf/177Hf of the St.Helena primary melt within 1 SD (Table 2).

Modeling results

The modeling results show that various, but limited com-binations of the modeling parameters can reproduce thetrace element abundances and isotopic ratios of the St.Helena primary melt (Table 5 and Fig. 16).The results indi-cate that the St. Helena primary melt can be produced by1^2% melting of the hybrid source (FOIB¼ 0·01^0·02)leaving olivine, orthopyroxene, clinopyroxene, and garnetin the residue. The hybrid source is explained by a mixtureof 82^92wt % model peridotite (XPeri¼ 0·82^0·92) and18^8wt % recycled crust-derived melt, which suggests


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that the lithology of the hybrid source is peridotitic.The mixing proportions of the model peridotite are con-sistent with those proposed by Chauvel et al. (1992).Chauvel et al. (1992) pointed out that the amount of perido-tite in the source of HIMU OIB should be higher than70% to reproduce the compatible element abundances ofHIMU OIB. The range of XPeri (0·82^0·92) also confirms

the conclusion in the previous section that peridotite withprimitive mantle or DMM compositions alone is unlikelyto be the source of the St. Helena primary melt.The model peridotite that was mixed with the

recycled crust-derived melt shows a depleted nature[XDMM¼ 0·7^1, (Rb/Nb)PM¼ 0·38^0·80]. Less depletedperidotite (XDMM50·7) is rejected, mainly because several

Fig. 15. Relationships between residual mineral abundance and XPeri. (a) Modeling results from the first calculation step. (b) Shaded areas rep-resent the residual mineral abundances calculated from pMELTS using major element compositions. Open symbols represent mineral abun-dances from the two sets of calculations that agree.


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elements (e.g. Rb and Pb) in the recycled crust be-come negative with a decrease in XDMM (Fig. 16f).The recycled crust-derived melt results from a highdegree of melting, 70^100% (FRC¼ 0·7^1). The DMM agevaries from 2·2 to 3·4 Ga, whereas the recycling age isrestricted to 1·2^2·8 Ga (Fig. 16e). The combination ofDMM age and recycling age is strongly constrainedby the Hf isotopic ratios, rather than the Pb, Sr, and Ndisotopic ratios.The calculated parameter values are geophysically rea-

sonable. For example, high FRC of 0·7^1 under near-solidusconditions for peridotite are supported by thermodynamiccalculations for adiabatic melting of recycled crustembedded in a peridotite-dominant mantle plume (PhippsMorgan, 2001). In addition, model OIB melts producedby a low degree of melting of a hybrid source(FOIB¼ 0·01^0·02) can segregate from the mantle, becausethey are beyond the melt segregation threshold estimatedfor natural peridotites (0·1^0·2 vol. % melt, Sundberget al., 2010).Recycled oceanic crust is denser than peridotite at

equivalent temperatures in the upper mantle (Yasudaet al., 1994; Pertermann & Hirschmann, 2003a). An excesstemperature of 100^2108C can sustain 10^18 wt % ofMORB-like eclogite in a mantle plume (Pertermann &Hirschmann, 2003a), if we adopt 3300 and 3440 kgm�3 asthe densities of peridotite and oceanic crust, respectively,and 4�10�58C�1 as the thermal expansion coefficient.This range of excess temperature is consistent with thatpreviously inferred for HIMU OIB in French Polynesia(Herzberg & Gazel, 2009).

Hybrid source compositions

The estimated hybrid source composition is shown inFig. 18e and listed in Table 5. The hybrid source shows abroadly similar element pattern to N-MORB (Fig. 18e),but element concentrations are 0·1^0·2 times theN-MORB values of Hofmann (1988). The hybrid sourcehas LREE/MREE (middle REE) ratios lower than PM[(La/Sm)PM¼ 0·57^0·64], whereas the MREE/HREEratio is broadly similar to that of PM [(Gd/Yb)PM¼0·98^1·18]. It should be noted that the low (Rb, Ba,Th, U)/Nb ratios in the source of HIMUOIB are not pro-duced only by recycled crust, but require a contributionfrom depleted peridotite.

Recycled crust isotopic compositions

The calculated recycled crust is allowed to have a widerange of present-day Pb isotopic ratios (206Pb/ 204Pb of21·7^73·9 and 208Pb/ 204Pb of 40·7^84·4), which is slightlyto significantly higher than those of St. Helena basalts(206Pb/ 204Pb of 20·7�0·3, 208Pb/ 204Pb of 40·0� 0·2)(Fig. 17). In contrast, the recycled crust has 87Sr/ 86Sr of0·7017^0·7042, 143Nd/144Nd of 0·51268^0·51285, and

176Hf/177Hf of 0·282572^0·28294. Overall, the Sr and Pbisotopic ratios in the recycled crust negatively correlate,whereas the Nd and Hf isotopic ratios positively correlate,with XDMM (Table 5).The estimated present-day 87Rb/ 86Sr, 147Sm/144Nd, and

176Lu/177Hf in the recycled crust (Table 5) fall within therange proposed in previous studies (Hauri & Hart, 1993;Stracke et al., 2003). In contrast, the 238U/ 204Pb and232Th/ 204Pb values obtained in this study extend therange of previous studies to higher values, although thereis some overlap.Previous studies have suggested that both recycled

oceanic crust and peridotite contribute to the genesis ofHIMU OIB. However, they neglected to consider theeffect of the peridotite, and assumed that the recycledoceanic crust had the same isotopic composition and in-compatible element ratios as present-day HIMU OIB.Our modeling results indicate that this approximation isan oversimplification, because XPeri is relatively large(0·82^0·92) and the contrast in element abundances is notsmall enough for the effect of peridotite to be ignored (e.g.ratios of element abundances in the recycled crust-derivedmelt to that in peridotite are 4^10 for Sr, 0·4^19 for Pb,13^27 for Nd). Thus, it is important to incorporate the in-fluence of the isotopic and element ratios of the peridotitein the genesis of HIMUOIB.

Recycled crust trace element compositions

The green shaded areas in Fig. 18a^e show the elementabundances and element ratios of the calculated recycledcrust. Also shown in Fig. 18a^d are compiled data formodern N-MORB glasses [(La/Sm)PM¼ 0·63^0·87],whose conservative element abundances fall within 15%deviation of the average N-MORB glass of Hofmann(1988). Modern N-MORB, which shows a wide range ofnon-conservative element abundances at given conserva-tive element abundances, is considered to reflect the origin-al composition of the recycled crust before seaflooralteration and slab dehydration.The calculated recycled crust is generally depleted in

Rb, Ba, La, Ce, Pb, and Sr compared with modernN-MORB, although the element abundances partly over-lap (e.g. Fig. 18a). The compositional difference betweenthe recycled crust and modern N-MORB can be explainedby a net effects of seafloor alteration and dehydrationduring subduction, as we discuss below.Two model dehydrated slab compositions are shown by

the hatched areas in Fig. 18a and b and Fig. 18c and d,which were calculated by modifying trace element behav-ior during slab dehydration. The dehydrated slab compos-ition in Fig. 18a and b was obtained using partitioncoefficients between mineral/fluid obtained from dehydra-tion experiments at 6GPa and 8008C (Kessel et al., 2005).Here the modal batch melting equation (Shaw, 1970) was


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Table 5: Calculated hybrid source and recycled crust compositions, and compiled modern N-MORB compositions

Modeling results Compiled N-MORB

XDMM¼ 1 XDMM¼ 0·9 XDMM¼ 0·8 XDMM¼ 0·7

Min. Max. Min. Max. Min. Max. Min. Max. Min. Max.

XPeri 0·82 0·92 0·82 0·92 0·83 0·92 0·84 0·87

FOIB 0·010 0·016 0·011 0·017 0·014 0·018 0·018 0·019

FRC 0·80 1·00 0·70 1·00 0·70 1·00 0·90 1·00

XOL 0·46 0·53 0·46 0·53 0·47 0·53 0·48 0·50

XOpx 0·17 0·23 0·16 0·22 0·16 0·22 0·17 0·18

XCpx 0·18 0·27 0·18 0·28 0·19 0·27 0·25 0·26

XGrt 0·06 0·10 0·06 0·10 0·06 0·10 0·08 0·09

DMM age (Ga) 2·2 3·4 2·2 3·4 2·2 3·4 2·4 3·4

Recycling age (Ga) 1·4 2·8 1·2 2·6 1·2 2·6 1·2 2·2

Hybrid source (ppm)

Rb 0·13 0·25 0·13 0·27 0·16 0·27 0·20 0·27 Rb 0·90 2·03

Ba 1·75 3·63 1·75 4·18 2·42 4·05 3·12 4·18 Ba 9·00 23·00

Th 0·03 0·06 0·03 0·07 0·04 0·07 0·05 0·07 Th 0·14 0·28

U 0·01 0·02 0·01 0·02 0·01 0·02 0·02 0·02 U 0·06 0·09

Nb 0·49 0·75 0·54 0·80 0·62 0·81 0·78 0·83 Nb 3·05 3·83

La 0·36 0·75 0·36 0·81 0·46 0·80 0·58 0·81 La 3·50 4·67

Ce 1·20 2·53 1·20 2·64 1·44 2·63 1·82 2·63 Ce 11·00 14·52

Pb 0·03 0·06 0·03 0·06 0·05 0·06 0·06 0·06 Pb 0·42 0·69

Pr 0·23 0·48 0·23 0·50 0·27 0·49 0·34 0·49 Pr 1·87 2·27

Sr 10·22 18·29 11·88 19·10 13·57 19·01 17·91 18·94 Sr 92 146

Nd 1·58 2·53 1·66 2·61 1·84 2·57 2·30 2·53 Nd 10·18 12·70

Sm 0·55 0·85 0·57 0·87 0·63 0·86 0·76 0·84 Sm 3·48 4·25

Zr 15·10 22·93 15·66 23·43 17·26 22·95 21·08 22·49 Zr 94 118

Hf 0·44 0·66 0·46 0·68 0·50 0·66 0·61 0·65 Hf 2·76 3·40

Eu 0·22 0·32 0·23 0·32 0·25 0·32 0·29 0·31 Eu 1·25 1·50

Gd 0·82 1·21 0·84 1·23 0·91 1·20 1·10 1·17 Gd 4·24 5·70

Dy 1·03 1·55 1·05 1·57 1·09 1·53 1·38 1·49 Dy 5·60 6·74

Y 6·08 9·18 6·20 9·28 6·30 9·05 8·07 8·83 Y 31 40

Ho 0·22 0·34 0·23 0·34 0·23 0·33 0·30 0·32 Ho 1·20 1·40

Er 0·66 1·03 0·67 1·04 0·68 1·02 0·89 0·99 Er 3·30 4·35

Yb 0·64 1·00 0·65 1·01 0·65 0·99 0·86 0·96 Yb 3·38 4·15

Lu 0·10 0·15 0·10 0·15 0·10 0·15 0·13 0·15 Lu 0·52 0·58

Recycled crust (ppm)

Rb 0·83 1·15 0·33 0·99 0·06 0·77 0·04 0·48

Ba 12·36 17·69 6·59 17·69 6·78 14·80 6·59 13·06

Th 0·21 0·32 0·18 0·30 0·17 0·29 0·17 0·27

U 0·07 0·10 0·06 0·10 0·06 0·10 0·06 0·10

La 1·86 3·34 1·75 3·37 1·75 3·35 1·93 3·32

Ce 6·95 12·00 6·56 12·32 6·56 12·24 7·02 11·77

Pb 0·12 0·27 0·02 0·20 0·02 0·12 0·02 0·05

Pr 1·33 2·29 1·29 2·37 1·29 2·37 1·38 2·25

Sr 31·73 66·69 35·94 65·06 36·93 61·27 53·55 56·96

Nd 10·40 11·84 10·46 12·20 10·55 12·20 10·90 11·67

Sm 3·07 3·70 3·30 3·79 3·31 3·76 3·42 3·67

(continued)


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adopted assuming 4wt % H2O loss (Schmidt & Poli,1998).The other dehydrated slab composition (Fig. 18c and d)was estimated using the element mobility when naturalamphibole is dehydrated at 5·5GPa and 9008C (Kogisoet al., 1997). It should be noted that the model dehydratedslab compositions represent minimum values for elementssuch as Rb and U, and occasionallyTh and Ba, consider-ing that these elements are also added during alteration(Staudigel et al., 1996; Alt & Teagle, 2003; Bach et al.,2003; Kelley et al., 2003). The extent of element depletionduring slab dehydration is different between the twoexperiments (Kogiso et al., 1997; Kessel et al., 2005). For ex-ample, loss of Th, U, Pb, La, Ce, and Nd, and fraction-ation of Nd^Sm, Th^U, Sr^Pb, Th^Pb, U^Pb are moresignificant in the experiments of Kogiso et al. (1997)(Fig. 18a^d). Thus, it is hard to estimate the ultimate com-position of the fully dehydrated slab reliably by simplyadopting the element mobility or partition coefficients ob-tained at specific P^T conditions. In this context, the

inverse modeling approach used in this study is more ef-fective in evaluating the oceanic crust-recycling hypoth-esis. Qualitatively, however, relative element mobility issimilar between the two experiments: Rb4Ba4Th4U,Rb4Sr, Pb4(Th, U), and Nd4Sm. The element frac-tionation expected during dehydration can qualitativelyexplain the trace element patterns (Fig. 18a and c) andthe change in trace element ratios in the recycledcrust (Fig. 18b and d). For example, recycled crust isdepleted in La, Ce, Pb, and Sr compared with modernN-MORB, and the extent of depletion in Th, Ba, and Rbincrease in order, both of which are well explained by slabdehydration. The recycled crust requires limited fraction-ation of Nd^Sm and Th^U, which is also consistent withthe experiments of Kessel et al. (2005) (Fig. 18a and b).The recycled crust tends to have higher Rb, Ba, and Rb/

Sr as XDMM increases (Fig. 16f and Table 5). In particular,recycled crust compositions calculated for XDMM4c. 0·9have maximum Rb contents of 0·9^1·2 ppm, which partly

Table 5: Continued

Modeling results

XDMM¼ 1 XDMM¼ 0·9 XDMM¼ 0·8 XDMM¼ 0·7

Min. Max. Min. Max. Min. Max. Min. Max.

Present-day P/D of recycled crust87Rb/86Sr 0·03958 0·09850 0·02244 0·05890 0·00383 0·04090 0·00217 0·02434147Sm/144Nd 0·17675 0·20071 0·17779 0·20218 0·17806 0·20437 0·17896 0·20167176Lu/177Hf 0·02809 0·02810 0·02809 0·02810 0·02809 0·02810 0·02809 0·02810238U/204Pb 16·8 37·1 20·4 178·3 29·3 238·2 68·8 230·7232Th/204Pb 58·5 117·9 70·0 541·9 99·9 699·5 218·4 676·9

Present-day isotopic ratios of recycled crust87Sr/86Sr 0·70306 0·70420 0·70276 0·70289 0·70173 0·70266 0·70174 0·70220143Nd/144Nd 0·51268 0·51279 0·51273 0·51281 0·51279 0·51283 0·51284 0·51285206Pb/204Pb 21·6 25·3 23·1 62·0 26·3 73·9 42·0 69·5208Pb/204Pb 40·7 44·0 42·1 74·5 45·2 84·4 58·9 83·3176Hf/177Hf 0·28257 0·28289 0·28258 0·28291 0·28263 0·28293 0·28269 0·28294

Present-day isotopic ratios of model peridotite87Sr/86Sr 0·70230 0·70293 0·70340 0·70376143Nd/144Nd 0·51325 0·51312 0·51302 0·51294206Pb/204Pb 17·7 17·7 17·7 17·7208Pb/204Pb 37·4 37·4 37·4 37·4176Hf/177Hf 0·28326 0·28318 0·28311 0·28305

Present-day isotopic ratios of model OIB87Sr/86Sr 0·70276 0·70289 0·70283 0·70291 0·70289 0·70302 0·70289 0·70302143Nd/144Nd 0·51287 0·51287 0·51287 0·51287 0·51287 0·51287 0·51287 0·51287206Pb/204Pb 20·56 20·96 20·56 20·97 20·55 20·97 20·66 20·93208Pb/204Pb 39·78 40·22 39·78 40·22 39·79 40·20 39·94 40·00176Hf/177Hf 0·28279 0·28297 0·28279 0·28297 0·28279 0·28297 0·28279 0·28297


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overlap with the range of modern MORB (0·9^2·0 ppm).Because slab dehydration predicts preferential Rb losscompared with Sr, Rb gain by seafloor alteration is essen-tial to explain recycled crust compositions that exhibit nosignificant Rb depletion. Consequently, if we consider thecombined element fractionation expected from seaflooralteration and slab dehydration, the compositional

difference between modern N-MORB and the calculatedrecycled oceanic crust can be explained. This indicatesthat the petrogenetic model shown in Fig. 14 describes apotential process to produce St. Helena primary meltswith a HIMU isotopic signature. Dehydrated oceaniccrust thus played an important role in producing the geo-chemistry of the St. Helena basalts.

Fig. 16. Combinations of model parameters that reproduce both the trace element abundances and isotopic ratios of the St. Helena basalts(a^e).With decreasing XDMM, calculated Rb and Pb abundances in the recycled crust decrease (f).


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Currently the model provides a wide range of potentialcompositions for the recycled crust component becausethe combinations of modeling parameters are still notwell constrained (Figs 17 and 18a^d). Among the free mod-eling parameters, XDMM strongly affects Rb, Ba, Pb andthe isotopic composition of the recycled crust. Thus, if wecan reliably estimate the degree of depletion of the modelperidotite, the potential range of recycled crust compos-itions may be further constrained.

Effect of subducted sediments

In an actual subduction system, sediment is subducted to-gether with the oceanic crust. Porter & White (2009) esti-mated the composition of the residual bulk slab subductedinto the deep mantle for eight representative modern arc

systems, based on a mass balance for trace element abun-dances. They assumed that the down-going bulk slab com-position is a mixture of sediment and oceanic crust (20%altered oceanic crust and 80% fresh oceanic crust).We tested whether such residual bulk slab compositions

are a likely component in the petrogenesis of St. Helenabasalts using the modeling scheme in Fig. 14. In this case,Nb, Zr, Hf, Eu,Y,Yb, and Lu in the average residual bulkslab (Porter & White, 2009) were substituted for therecycled crust compositions in the first calculation step(Appendix Fig. A1). The fourth calculation step wasskipped here, because there are large uncertainties in theestimation of the isotopic composition of the subductedsediment. The gray field in Fig. 18f shows the calculatednon-conservative element abundances in the recycled crust.

Fig. 17. Calculated isotopic compositions of the recycled crust, model OIB, and model peridotite.The isotopic ratios of recycled crust are calcu-lated to reproduce the isotopic ratios of the St. Helena basalts based on the petrogenetic model shown in Fig. 14. For comparison, published iso-topic compositions of St. Helena basalts, MORB and OIB are also shown.


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Fig. 18. (a^d) Comparison of the calculated recycled crust compositions with modern N-MORB and model dehydrated N-MORB. ModernN-MORB compositions are from Van Wagoner & Leybourne (1991), Goldstein et al. (1993), Harpp (1995), Reynolds (1995), Meyer & Bryan(1996), Regelous et al. (1999), le Roux (2000), Mahoney et al. (2002), Sims et al. (2002, 2003), and le Roux et al. (2006). Model dehydratedN-MORB was calculated using the compiled modern MORB compositions (Table 5) and element behavior during slab dehydration. In (a)and (b), the model dehydrated N-MORB composition was calculated using mineral^fluid partition coefficients at 6GPa and 8008C (Kesselet al., 2005), assuming 4 wt % H2O loss. In (c) and (d), the model dehydrated N-MORB composition was calculated using the element mobilityin natural amphibolite during dehydration at 5·5GPa and 9008C (Kogiso et al., 1997). Trace element abundances are normalized to theN-MORB values of Hofmann (1988). (e) Calculated recycled crust and hybrid source compositions normalized to PM values. For comparison,the compositions of the St. Helena primary melt, DMM (Workman & Hart, 2005), and N-MORB (Hofmann, 1988) are also shown.(f) Calculated recycled crust composition assuming that the average residual bulk slab (Porter & White, 2009) contributed to the genesis ofthe St. Helena basalts. This calculation was performed by modifying the modeling scheme in Fig. 14 (see text). Thus, the calculated recycledcrust compositions are different between (f) and (a)^(e). Prominent compositional differences in Rb, Ba, U, and Pb between the residual bulkslab and the calculated recycled crust suggest that subducted sediment is not involved in the genesis of St. Helena basalts.


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The calculated recycled crust compositions are distinctfrom the average element abundances in the residual bulkslab, particularly those of Rb, Ba, U, and Pb, suggestingthat the residual bulk slab is not related to the genesis ofthe St. Helena basalts (Fig. 18f). This result is consistentwith the conclusion of Porter & White (2009) based on Sr,Nd, and Pb isotopic compositions. The residual bulk slabcompositions estimated by Porter & White (2009) are sig-nificantly enriched in Rb, Ba, U and Pb compared withfresh N-MORB, mainly owing to the incorporation of sedi-ment in the bulk slab. Thus, this result suggests thatrecycled sediment is not involved in the genesis of the St.Helena basalts.

Model limitations and other potential models

As mentioned above, we have made the following assump-tions in developing our geochemical model: (1) the sub-ducted oceanic crust that contributed to the genesis of theSt. Helena primary basalts has N-MORB composition; (2)conservative elements are not mobile during seafloor alter-ation and slab dehydration; (3) the source mantle of theSt. Helena basalts is approximated by a mixture ofslab-derived melt and peridotite. Thus our petrogeneticmodel is a straightforward oceanic crust-recycling model.It should be noted that a wide variety of petrogenetic scen-arios are possible in the formation of OIB, depending onthe various combinations of processes operating in theoceanic crust, prior to and during subduction, and duringsubsequent mantle convection. For example, the role ofother materials that make up the oceanic crust (e.g. sedi-ment, cumulate gabbros), and slab melting during sub-duction and in the lower mantle, on the genesis of OIBshould be examined. Moreover, recycled lithosphericmantle peridotite metasomatized by low-degreemelts couldalso be a potential source component for OIB (Niu & O’Hara, 2003; Pilet et al., 2005). These potential petrogeneticprocesses may be quantitatively examined using a similarinverse approach to that proposed here, if we describe suit-able petrogenetic scenarios for eachpotentialmodel.

CONCLUSIONSSt. Helena alkali basalts share common geochemical char-acteristics with HIMU OIB from French Polynesia. Theirpetrographic and geochemical characteristics indicatethat the bulk-rock compositions of the St. Helena basaltsrepresent a mixture of more evolved melt and variousphenocryst phases. Careful examination of the petrologyand geochemistry indicates that the key factors that pro-duced compositional variations in the alkali basalts arethe extent of fractional crystallization and the accumula-tion of olivine and clinopyroxene phenocrysts.We estimated the primary magma composition by calcu-

lating the reverse path of fractional crystallization basedon the reconstructed order of crystallization and relative

amount of fractionated phases. This back calculation indi-cates that the St. Helena alkali basalts were derived froma common primary magma with 14^20wt % MgO.The St. Helena primary melt cannot be explained by the

partial melting of fertile peridotite (pyrolite), depletedMORB-source mantle, garnet pyroxenite, or hornblendite.However, partial melts of metasomatized Fe-rich peridotiteand of the reaction products between partial melts ofMORB-like eclogite and peridotite show broadly similarmajor element compositions to the St. Helena primarymelt. This suggests that a second component is required toproduce the St. Helena basalts, in addition to peridotite.Ancient subducted oceanic crust is the most likely sec-

ondary component. Thus, we tested whether an oceaniccrust-recycling model is likely for the genesis of the St.Helena basalts. Geochemical modeling using major elem-ent, trace element and isotopic compositions indicates thatthe St. Helena primary melt can be produced by 1^2%melting of a hybrid source leaving olivine, orthopyroxene,clinopyroxene, and garnet in the residue. The hybridsource is explained by a mixture of 82^92% model perido-tite and 18^8% recycled crust-derived melt, which suggeststhat the lithology of the hybrid source is peridotitic. Themodel peridotite that was hybridized with the recycledcrust-derived melt is depleted in nature.The geochemical model developed here calculates the

trace element abundances in and present-day isotopic com-position of the recycled crust that is required to reproducethe geochemistry of the St. Helena basalts, withoutmaking any assumptions regarding the behavior offluid-mobile elements during seafloor alteration and slabdehydration. The calculated recycled crust has a widerange of present-day Pb isotopic ratios (206Pb/ 204Pb of21·7^79·3 and 208Pb/ 204Pb of 40·8^89·3), 87Sr/ 86Sr(0·7018^0·7028), 143Nd/144Nd (0·51274^0·51285), and176Hf/177Hf (0·28262^0·28293) after a residence time of1·2^2·8 Gyr. It is depleted in Rb, Ba, Pb, and Sr comparedwith modern N-MORB. In contrast, for Th, U, Sm, andNd, the extent of chemical modification is restricted to therange of compositional variations in modern N-MORB.These geochemical characteristics of the recycled crustcan be qualitatively explained by element behavior duringseafloor alteration and slab dehydration. It should benoted that subducted sediment has to be excluded to ex-plain the low (Rb, Ba, U)/Nb and Ce/Pb ratios of the St.Helena basalts. These results suggest that oceanic crustrecycling is one of the most likely processes capable of pro-ducing HIMUOIB.

ACKNOWLEDGEMENTSWe thank Ian Baker, Hiroshi Shukuno, Ken Tani, andYoshiyuki Tamura for help and discussions. Constructivereviews by Catherine Chauvel, Andreas Stracke,


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Se¤ bastien Pilet and Editor Marjorie Wilson improved thequality of this paper and were much appreciated.

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

REFERENCESAbdel-Monem, A. & Gast, P. W. (1967). Age of volcanism on St.

Helena. Earth and Planetary Science Letters 2, 415^418.Albare' de, F. & Tamagnan, V. (1988). Modelling the recent geochem-

ical evolution of the Piton de la Fournaise volcano, ReunionIsland, 1931^1986. Journal of Petrology 29, 997^1030.

Alt, J. C. & Teagle, D. A. H. (2003). Hydrothermal alteration of upperoceanic crust formed at a fast-spreading ridge: mineral, chemical,and isotopic evidence from ODP Site 801. Chemical Geology 201,191^211.

Arai, S. & Fujii, T. (1979). Petrology of ultramafic rocks from Site 395.In: Melson,W. G. & Rabinowicz, P. D. et al. (eds) Initial Reports ofthe Deep Sea Drilling Project, 45. Washington, DC: US GovernmentPrinting Office, pp. 587^594.

Arai, S. & Matsukage, K. (1996). Petrology of the gabbro^troctolite^peridotite complex from Hess deep, equatorial Pacific:Implications for mentle^melt interaction within the oceanic litho-sphere. In: Me¤ vel, C., Gillis, K. M., Allan, J. F. & Meyer, P. S.(eds) Proceedings of the Ocean Drilling Program, Scientific Results,Vol 147. College Station,TX: Ocean Drilling Program, pp. 135^155.

Bach,W., Peucker-Ehrenbrink, B., Hart, S. R. & Blusztajn, J. S. (2003).Geochemistry of hydrothermally altered oceanic crust: DSDP/ODP Hole 504BçImplications for seawater^crust exchange bud-gets and Sr- and Pb-isotopic evolution of the mantle. Geochemistry,Geophysics, Geosystems 4, 8904.

Baker, I. (1968a). The geology of Saint Helena Island, South Atlantic.PhD thesis. London University, 369 p.

Baker, I. (1968b). Intermediate oceanic volcanic rocks and the ‘DalyGap’. Earth and Planetary Science Letters 4, 103^106.

Baker, I. (1969). Petrology of the volcanic rocks of Saint Helena Island,South Atlantlc. Geological Society of America Bulletin 80, 1283^1310.

Baker, I., Gale, N. H. & Simons, J. (1967). Geochronology of theSt. Helena volcanoes. Nature 215, 1451^1456.

Beier, C., Stracke, A. & Haase, K. M. (2007). The peculiar geochem-ical signatures of Sa‹ o Miguel (Azores) lavas: Metasomatised orrecycled mantle sources? Earth and Planetary Science Letters 259,186^199.

Brenan, J. M., Shaw, H. F., Ryerson, F. J. & Phinney, D. L. (1995).Mineral^aqueous fluid partitioning of trace elements at 9008Cand 2·0GPa: Constraints on the trace element chemistry ofmantle and deep crustal fluids. Geochimica et Cosmochimica Acta 59,3331^3350.

Brunelli, D., Seyler, M., Cipriani, A., Ottolini, L. & Bonatti, E.(2006). Discontinuous melt extraction and weak refertilization ofmantle peridotites at theVema Lithospheric Section (Mid-AtlanticRidge). Journal of Petrology 47, 745^771.

Cande, S. C. & Kent, D. V. (1992). A new geomagnetic polarity timescale for the Late Cretaceous Cenozoic. Journal of Geophysical

Research 97, 13917^13951.Cantagrel, J. M., Jamond, C. & Lassere, M. (1978). Le magmatism

alcalin de la ligne du Cameroun au Tertiaire infe¤ rieur: donne¤ esge¤ ochronologiques K/Ar. Comptes Rendus Sommaire de la Socie¤ te¤

Ge¤ ologique de France 6, 300^303.

Carmichael, L. S. E. (1967). The iron^titanium oxides of salic volcanicrocks and their associated ferromagnesian silicates. Contributions toMineralogy and Petrology 14, 36^64.

Chaffey, D. J., Cliff, R. A. & Wilson, B. M. (1989). Characterization ofthe St Helena magma source. In: Saunders, A. D. & Norry, M. J.(eds) Magmatism in the Ocean Basins. Geological Society, London, Special

Publications 42, 257^276.Champion, K. P., Taylor, J. C. & Whittem, R. N. (1966). Rapid X-ray

fluorescence determination of traces of strontium in samples of bio-logical and geological origin. Analytical Chemistry 38, 109^112.

Chang, Q., Shibata, T., Shinotsuka, K., Yoshikawa, M. & Tatsumi, Y.(2003). Precise determination of trace elements in geological stand-ard rocks using inductively coupled plasma mass spectrometry(ICP-MS). Frontier Research on Earth Evolution 1, 357^362.

Chauvel, C., Hofmann, A. W. & Vidal, P. (1992). HIMU-EM: TheFrench Polynesian connection. Earth and Planetary Science Letters

110, 99^119.Clague, D. A. & Denlinger, R. P. (1994). Role of olivine cumulates in

destabilizing the flanks of Hawaiian volcanoes. Bulletin of

Volcanology 56, 425^434.Clarke, D. B., McCuish, K. L., Vernon, R. H., Maksaev, V. &

Miller, B. V. (2002). The Port Mouton Shear Zone: intersection ofa regional fault with a crystallizing granitoid pluton. Lithos 61,141^159.

Dasgupta, R., Hirschmann, M. M. & Smith, N. D. (2007). Partialmelting experiments of peridotite þ CO2 at 3GPa and genesis ofalkalic ocean island basalts. Journal of Petrology 48, 2093^2124.

D’Orazio, M., Armienti, P. & Cerretini, S. (1998). Phenocryst/matrixtrace-element partition coefficients for hawaiite^trachyte lavasfrom the Ellittico volcanic sequence (Mt. Etna, Sicily, Italy).Mineralogy and Petrology 64, 65^88.

Ellam, R. M. (1992). Lithospheric thickness as a control on basalt geo-chemistry. Geology 20, 153^156.

Fitton, J. G. & Dunlop, H. M. (1985).The Cameroon line,West Africa,and its bearing on the origin of oceanic and continental alkalibasalt. Earth and Planetary Science Letters 72, 23^38.

Fujii, T. (1990). Petrology of peridotites from Hole 670A, Leg 109.In: Detrick, R., Honnorez, J., Bryan, W. B. & Juteau, T. (eds)Proceedings of the Ocean Drilling Program, Scientific Results, Vol 106/109.College Station,TX: Ocean Drilling Program, pp. 19^25.

Galer, S. J. G. & O’Nions, R. K. (1985). Residence time of thorium, ur-anium and lead in the mantle with implications for mantle convec-tion. Nature 316, 778^782.

Garcia, M. O. (1996). Petrography and olivine and glass chemistry oflavas from the Hawaii Scientific Drilling Project. Journal of

Geophysical Research 101, 11701^11713.Ghiorso, M. S., Hirschmann, M. M., Reiners, P. W. & Kress, V. C.

(2002). The pMELTS: A revision of MELTS for improved calcula-tion of phase relations and major element partitioning related topartial melting of the mantle to 3GPa. Geochemistry, Geophysics,

Geosystems 3, doi: 10.1029/2001GC000217.Goldstein, J., Newbury, D. E., Joy, D. C., Lyman, C. E., Echlin, P.,

Lifshin, E., Sawyer, L. & Michael, J. R. (2003). Scanning Electron

Microscopy and X-ray Microanalysis. Berlin: Springer.Goldstein, S. J., Murrell, M. T. & Williams, R. W. (1993). 231Pa and

230Th chronology of mid-ocean ridge basalts. Earth and Planetary

Science Letters 115, 151^159.Goto, A. & Tatsumi, Y. (1996). Quantitative analyses of rock samples

by X-ray fluorescence spectrography (II). RigakuJournal 13, 20^38.Halliday, A. N., Lee, D.-C., Tommasini, S., Davies, G. R., Paslick, C.

R., Godfrey Fitton, J. & James, D. E. (1995). Incompatible traceelements in OIB and MORB and source enrichment in thesub-oceanic mantle. Earth and Planetary Science Letters 133, 379^395.


830

Dow



Hanyu,T. & Kaneoka, I. (1997).The uniform and low 3He/4He ratiosof HIMU basalts as evidence for their origin as recycled materials.Nature 390, 273^276.

Hanyu,T.,Tatsumi,Y., Senda,R.,Miyazaki,T.,Chang,Q.,Hirahara,Y.,Takahashi,T., Kawabata, H., Suzuki, K., Kimura, J.-I. & Nakai, S(2011). Geochemical characteristics and origin of the HIMU reser-voir: A possible mantle plume source in the lower mantle.Geochemistry,Geophysics,Geosystems doi:10.1029/2010GC003252.

Harpp, K. S. (1995). Geochemical variations in a single basaltic flowat 98300N on the East Pacific Rise. In: Batiza, R., Storms, M.A. &Allan, J.F. (eds) Proceedings of the Ocean Drilling Program, Scientific

Results, 142. College Station, TX: Ocean Drilling Program,pp. 9^22.

Hart, S. & Gaetani, G. (2006). Mantle Pb paradoxes: the sulfide solu-tion. Contributions to Mineralogy and Petrology 152, 295^308.

Hart, S. R. (1988). Heterogeneous mantle domains: signatures, genesisand mixing chronologies. Earth and Planetary Science Letters 90,273^296.

Hart, S. R., Hauri, E. H., Oschmann, L. A. & Whitehead, J. A.(1992). Mantle plumes and entrainment: Isotopic evidence. Nature256, 517^520.

Hauri, E. H. & Hart, S. R. (1993). Re^Os isotope systematics ofHIMU and EMII oceanic island basalts from the South PacificOcean. Earth and Planetary Science Letters 114, 353^371.

Hellebrand, E., Snow, J. E., Hoppe, P. & Hofmann, A. W. (2002).Garnet-field melting and late-stage refertilization in ‘residual’abys-sal peridotites from the Central Indian Ridge. Journal of Petrology43, 2305^2338.

Herzberg, C. & Gazel, E. (2009). Petrological evidence for secularcooling in mantle plumes. Nature 458, 619^622.

Herzberg, C., Raterron, P. & Zhang, J. (2000). New experimentalobservations on the anhydrous solidus for peridotite KLB-1.Geochemistry, Geophysics, Geosystems 1, doi: 10.1029/2000GC000089.

Hilton, D. R., Macpherson, C. G. & Elliott, T. R. (2000). Helium iso-tope ratios in mafic phenocrysts and geothermal fluids from LaPalma, the Canary Islands (Spain): Implications for HIMUmantle sources. Geochimica et Cosmochimica Acta 64, 2119^2132.

Hirose, K. & Kushiro, I. (1993). Partial melting of dry peridotites athigh pressures: Determination of compositions of melts segregatedfrom peridotite using aggregates of diamond. Earth and Planetary

Science Letters 114, 477^489.Hirschmann, M. M. (2000). Mantle solidus: Experimental constraints

and the effects of peridotite composition. Geochemistry, Geophysics,Geosystems 1, doi: 10.1029/2000GC000070.

Hirschmann, M. M., Kogiso,T., Baker, M. B. & Stolper, E. M. (2003).Alkalic magmas generated by partial melting of garnet pyroxenite.Geology 31, 481^484.

Hofmann, A.W. (1988). Chemical differentiation of the Earth: the re-lationship between mantle, continental crust, and oceanic crust.Earth and Planetary Science Letters 90, 297^314.

Hofmann, A.W. & White,W. M. (1982). Mantle plumes from ancientoceanic crust. Earth and Planetary Science Letters 57, 421^436.

Humphreys, E. R. & Niu, Y. (2009). On the composition of oceanisland basalts (OIB): The effects of lithospheric thickness variationand mantle metasomatism. Lithos 112, 118^136.

Imai, N., Terashima, S., Itoh, S. & Ando, A. (1995). 1994 compilationof analytical data for minor and trace elements in seventeen GSJgeochemical reference samples, ‘Igneous Rock Series’. GeostandardsNewsletter 19, 135^213.

Ito, G. & Mahoney, J. J. (2005). Flow and melting of a heterogeneousmantle: 1. Method and importance to the geochemistry of oceanisland and mid-ocean ridge basalts. Earth and Planetary Science

Letters 230, 29^46.

Jacquemin, H., Sheppard, S. M. F. & Vidal, P. (1982). Isotopic geo-chemistry (O, Sr, Pb) of the Golda Zuelva andMboutou anorogeniccomplexes, North Cameroun: mantle origin with evidence forcrustal contamination. Earth and Planetary Science Letters 61, 97^111.

Jochum, K. P., Nohl, U., Herwig, K., Lammel, E., Stoll, B. &Hofmann, A. W. (2005). GeoReM: A new geochemical databasefor reference materials and isotopic standards. Geostandards and

Geoanalytical Research 29, 333^338.Johnson, K.T. M. (1998). Experimental determination of partition co-

efficients for rare earth and high-field-strength elements betweenclinopyroxene, garnet, and basaltic melt at high pressures.Contributions to Mineralogy and Petrology 133, 60^68.

Juteau, T., Berger, E. & Cannat, M. (1990). Serpentinized, residualmantle peridotites from the M.A.R. median valley, ODP Hole67,A (218100N, 45800W, Leg 109): Primary mineralogy and geother-mometry. In: Detrick, R., Honnorez, J., Bryan,W. B. & Juteau, T.(eds) Proceedings of the Ocean Drilling Program, Scientific Results,Vol. 106/109. College Station, TX: Ocean Drilling Program,pp. 27^45.

Kelley, K. A., Plank, T., Ludden, J. & Staudigel, H. (2003).Composition of altered oceanic crust at ODP Sites 801 and 1149.Geochemistry, Geophysics, Geosystems 4, 8910.

Keppler, H. (1996). Constraints from partitioning experiments on thecomposition of subduction-zone fluids. Nature 380, 237^240.

Keshav, S., Gudfinnsson, G. H., Sen, G. & Fei, Y. (2004).High-pressure melting experiments on garnet clinopyroxenite andthe alkalic to tholeiitic transition in ocean-island basalts. Earth andPlanetary Science Letters 223, 365^379.

Kessel, R., Schmidt, M.W., Ulmer, P. & Pettke,T. (2005). Trace elem-ent signature of subduction-zone fluids, melts and supercritical li-quids at 120^180 km depth. Nature 437, 724^727.

Kinzler, R. J., Grove, T. L. & Recca, S. I. (1990). An experimentalstudy on the effect of temperature and melt composition on the par-titioning of nickel between olivine and silicate melt. Geochimica et

Cosmochimica Acta 54, 1255^1265.Klemme, S., Blundy, J. D. & Wood, B. J. (2002). Experimental con-

straints on major and trace element partitioning during partialmelting of eclogite. Geochimica et Cosmochimica Acta 66, 3109^3123.

Klemme, S., Prowatke, S., Hametner, K. & Gunther, D. (2005).Partitioning of trace elements between rutile and silicate melts:Implications for subduction zones. Geochimica et Cosmochimica Acta

69, 2361^2371.Kogiso, T., Tatsumi, Y. & Nakano, S. (1997). Trace element transport

during dehydration processes in the subducted oceanic crust: 1.Experiments and implications for the origin of ocean island bas-alts. Earth and Planetary Science Letters 148, 193^205.

Kogiso,T., Hirose, K. & Takahashi, E. (1998). Melting experiments onhomogeneous mixtures of peridotite and basalt: application to thegenesis of ocean island basalts. Earth and Planetary Science Letters

162, 45^61.Kogiso, T., Hirschmann, M. M. & Frost, D. J. (2003). High-pressure

partial melting of garnet pyroxenite: possible mafic lithologies inthe source of ocean island basalts. Earth and Planetary Science Letters216, 603^617.

Kogiso, T., Hirschmann, M. M. & Pertermann, M. (2004).High-pressure partial melting of mafic lithologies in the mantle.Journal of Petrology 45, 2407^2422.

Komor, S. C., Grove, T. L. & Rejean, H. (1990). Abyssal peridotitesfrom ODP Hole 670A (218100N, 458020W): Residues of mantle melt-ing exposed by non-constructive axial divergence. In: Detrick, R.,Honnorez, J., Bryan, W. B. & Juteau, T. (eds) Proceedings of

the Ocean Drilling Program, Scientific Results, Vol. 106/109. CollegeStation, TX: Ocean Drilling Program.


831

Dow



Kushiro, I. (1996). Partial melting of a fertile mantle peridotiteat high pressures: An experimental study using aggregates ofdiamond. In: Basu, A. & Hart, S. (eds). Earth Processes: Reading

Isotopic Code, American Geophysical Union, Geophysical Monograph 95,109^122.

Lawver, L. A., Gahagan, L. M. & Dalziel, I.W. D. (1999). A tight fitçEarly Mesozoic Gondwana, a plate reconstruction perspective.Memoirs of the National Institute of Polar Research, Special Issue 53,214^229.

Le Bas, M. J., Maitre, R.W. L., Streckeisen, A. & Zanettin, B. (1986).A chemical classification of volcanic rocks based on the totalalkali^silica diagram. Journal of Petrology 27, 745^750.

le Roux, P. J. (2000). The geochemistry of selected mid-oceanridge basalts from the southern Mid-Atlantic Ridge (40 to 558S),PhD thesis, University of CapeTown, South Africa.

le Roux, P. J., Shirey, S. B., Hauri, E. H., Perfit, M. R. & Bender, J. F.(2006). The effects of variable sources, processes and contaminantson the composition of northern EPR MORB (8^108N and 12^148N): Evidence from volatiles (H2O, CO2, S) and halogens (F,Cl). Earth and Planetary Science Letters 251, 209^231.

Libourel, G. (1999). Systematics of calcium partitioning between oliv-ine and silicate melt: implications for melt structure and calciumcontent of magmatic olivines. Contributions to Mineralogy and

Petrology 136, 63^80.Macdonald, G. A. & Katsura, T. (1964). Chemical composition of

Hawaiian lavas 1. Journal of Petrology 5, 82^133.Mahoney, J. J., Graham, D. W., Christie, D. M., Johnson, K. T. M.,

Hall, L. S. & Vonderhaar, D. L. (2002). Between a hotspot and acold spot: isotopic variation in the Southeast Indian Ridge as-thenosphere, 868E^1188E. Journal of Petrology 43, 1155^1176.

Marzoli, A., Piccirillo, E. M., Renne, P. R., Bellieni, G., Iacumin, M.,Nyobe, J. B. & Tongwa, A. T. (2000). The CameroonVolcanic Linerevisited: petrogenesis of continental basaltic magmas from litho-spheric and asthenospheric mantle sources. Journal of Petrology 41,87^109.

McDonough,W. F. & Sun, S. s. (1995). The composition of the Earth.Chemical Geology 120, 223^253.

McKenzie, D. & O’Nions, R. K. (1998). Melt production beneathoceanic islands. Physics of the Earth and Planetary Interiors 107,143^182.

Meyer, P. S. & Bryan,W. B. (1996). Petrology of basaltic glasses fromthe TAG Segment: Implications for a deep hydrothermal heatsource. Geophysical Research Letters 23, 3435^3438.

Montelli, R., Nolet, G., Dahlen, F. A., Masters, G., Engdahl, E. R. &Hung, S.-H. (2004). Finite-frequency tomography reveals a varietyof plumes in the mantle. Science 303, 338^343.

Niida, K. (1997). Mineralogy of Mark peridotites: Replacementthrough magma channeling examined from Hole 920D, Markarea. In: Karso, J. A., Cannat, M., Miller, D. J. & Eithon, D.(eds) Proceedings of the Ocean Drilling Program, Scientific Results,Vol. 153. College Station, TX: Ocean Drilling Program,pp. 265^275.

Niu,Y. & O’Hara, M. J. (2003). Origin of ocean island basalts: A newperspective from petrology, geochemistry, and mineral physics con-sideration. Journal of Geophysical Research 108, 2209, doi: 10.1029/2002JB002048.

O’Connor, J. M. & Duncan, R. A. (1990). Evolution of the WalvisRidge^Rio Grande Rise hot spot-system: Implications for Africanand South American plate motions over plumes. Journal of

Geophysical Research 95, 17475^17502.O’Connor, J. M. & le Roex, A. P. (1992). South Atlantic hot spot^

plume systems: 1. Distribution of volcanism in time and space.Earth and Planetary Science Letters 113, 343^364.

O’Hara, M. J. (1968). The bearing of phase equilibria studies on theorigin and evolution of igneous rocks. Earth-Science Reviews 4,69^133.

Pertermann, M. & Hirschmann, M. M. (2003a). Anhydrous partialmelting experiments on MORB-like eclogite: Phase relations,phase compositions and mineral^melt partitioning of major elem-ents at 2^3GPa. Journal of Petrology 44, 2173^2201.

Pertermann, M. & Hirschmann, M. M. (2003b). Partial melting ex-periments on a MORB-like pyroxenite between 2 and 3GPa:Constraints on the presence of pyroxenite in basalt source regionsfrom solidus location and melting rate. Journal of Geophysical

ResearchçSolid Earth 108, doi: 10.1029/2000JB000118.Phipps Morgan, J. (2001).Thermodynamics of pressure release melting

of a veined plum pudding mantle. Geochemistry, Geophysics, Geosystems2,1001, doi:10.1029/2000GC000049.

Pilet, S., Hernandez, J., Sylvester, P. & Poujol, M. (2005). The meta-somatic alternative for ocean island basalt chemical heterogeneity.Earth and Planetary Science Letters 236, 148^166.

Pilet, S. B., Baker, M. B. & Stolper, E. M. (2008). Metasomatizedlithosphere and the origin of alkaline lavas. Science 320, 916^919.

Porter, K. A. & White, W. M. (2009). Deep mantle subduction flux.Geochemistry, Geophysics, Geosystems 10, Q12016.

Prytulak, J. & Elliott,T. (2007). TiO2 enrichment in ocean island bas-alts. Earth and Planetary Science Letters 263, 388^403.

Regelous, M., Niu, Y., Wendt, J. I., Batiza, R., Greig, A. &Collerson, K. D. (1999). Variations in the geochemistry of magma-tism on the East Pacific Rise at 108300N since 800 ka. Earth and

Planetary Science Letters 168, 45^63.Reisberg, L., Zindler, A., Marcantonio, F., White, W., Wyman, D. &

Weaver, B. (1993). Os isotope systematics in ocean island basalts.Earth and Planetary Science Letters 120, 149^167.

Reynolds, J. (1995). Segment scale systematics of mid-ocean ridgemagmatism and geochemistry. Ph.D. Thesis, NewYork: ColumbiaUniversity.

Salters, V. J. M. & Stracke, A. (2004). Composition of thedepleted mantle. Geochemistry, Geophysics, Geosystems 5, doi: 10.1029/2003GC000597.

Salters, V. J., Longhi, J. E. & Bizimis, M. (2002). Near mantle solidustrace element partitioning at pressures up to 3·4GPa. Geochemistry,Geophysics, Geosystems 3, doi: 10.1029/2001GC000148.

Schmidt, M.W. & Poli, S. (1998). Experimentally based water budgetsfor dehydrating slabs and consequences for arc magma generation.Earth and Planetary Science Letters 163, 361^379.

Sears, J.W., St. George, G. M. & Winne, J. C. (2005). Continental riftsystems and anorogenic magmatism. Lithos 80, 147^154.

Sekine, T. & Wyllie, P. J. (1983). Experimental simulation of mantlehybridization in subduction zones. Journal of Geology 91, 511^528.

Sen, C. & Dunn, T. (1995). Experimental modal metasomatism of aspinel lherzolite and the production of amphibole-bearing perido-tite. Contributions to Mineralogy and Petrology 119, 422^432.

Shaw, D. M. (1970). Trace element fractionation during anatexis.Geochimica et Cosmochimica Acta 34, 237^243.

Shukuno, H. (2003). Quantitative analysis of rock-forming mineralsand volcanic glasses by electron probe microanalyser. Frontier

Research on Earth Evolution 1, 129^136.Sims, K. W. W., Goldstein, S. J., Blichert-Toft, J., Perfit, M. R.,

Kelemen, P., Fornari, D. J., Michael, P., Murrell, M. T., Hart, S.R., DePaolo, D. J., Layne, G., Ball, L., Jull, M. & Bender, J.(2002). Chemical and isotopic constraints on the generation andtransport of magma beneath the East Pacific Rise. Geochimica et

Cosmochimica Acta 66, 3481^3504.Sims, K. W. W., Blichert-Toft, J., Fornari, D. J., Perfit, M. R.,

Goldstein, S. J., Johnson, P., DePaolo, D. J., Hart, S. R.,


832

Dow



Murrell, M. T., Michael, P. J., Layne, G. D. & Ball, L. A. (2003).Aberrant youth: Chemical and isotopic constraints on the originof off-axis lavas from the East Pacific Rise, 98^108N. Geochemistry,Geophysics, Geosystems 4, 8621.

Sinton, J. M. (1979). Petrology of (Alpine-type) peridotites fromSite 395, DSDP Leg 45. In: Melson, W.G. & Rabinowitz, P.D.(eds) Initial Reports of the Deep Sea Drilling Project, 45. Washington,DC: US Government Printing Office, pp. 595^601.

Smith, D., Griffin,W. L. & Ryan, C. G. (1993). Compositional evolu-tion of high-temperature sheared lherzolite PHN 1611. Geochimica etCosmochimica Acta 57, 605^613.

Sobolev, A. V., Hofmann, A. W., Sobolev, S. V. & Nikogosian, I. K.(2005). An olivine-free mantle source of Hawaiian shield basalts.Nature 434, 590^597.

Staudigel, H., Plank, T., White, W. & Schmincke, H.-U. (1996).Geochemical fluxes during seafloor alteration of the basaltic upperoceanic crust: DSDP Sites 417 and 418 (overview). In: Be¤ bout, G. E.,Scholl, D.W., Kirby, S. H. & Platt, J. P. (eds) Subduction:Top to Bottom.AmericanGeophysicalUnion,GeophysicalMonograph 96,19^38.

Stracke, A. & Bourdon, B. (2009). The importance of melt extractionfor tracing mantle heterogeneity. Geochimica et Cosmochimica Acta 73,218^238.

Stracke, A., Bizimis, M. & Salters, V. J. (2003). Recycling oceaniccrust: Quantitative constraints. Geochemistry, Geophysics, Geosystems 4,doi:10.1029/2001GC000223.

Stracke, A., Hofmann, A. W. & Hart, S. R. (2005). FOZO, HIMU,and the rest of the mantle zoo. Geochemistry, Geophysics, Geosystems 6,doi: 10.1029/2004GC000824.

Sun, S. s. &McDonough,W. F. (1989). Chemical and isotopic systematicof oceanic basalts: implications for mantle composition and pro-cesses. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the

OceanBasins. Geological Society, London, Special Publications 42,313^345.Sundberg, M., Hirth, G. & Kelemen, P. B. (2010). Trapped melt in the

Josephine Peridotite: implications for permeability and melt ex-traction in the upper mantle. Journal of Petrology 51, 185^200.

Tani, K., Kawabata, H., Chang, Q., Sato, K. & Tatsumi, Y. (2006).Quantitative analyses of silicate rock major and trace elementsby X-ray fluorescence spectrometer: Evaluation of analyticalprecision and sample preparation. Frontier Research on Earth

Evolution 2, 1^8.Tatsumi, Y. (2000). Continental crust formation by crustal delamin-

ation in subduction zones and complementary accumulation ofthe enriched mantle 1 component in the mantle. Geochemistry,

Geophysics, Geosystems 1, doi: 10.1029/2000GC000094.Taura, H., Yurimoto, H., Kato, T. & Sueno, S. (2001). Trace element

partitioning between silicate perovskites and ultracalcic melt.Physics of the Earth and Planetary Interiors 124, 25^32.

Thy, P. (1995). Experimental constraints on the evolution of transition-al and mildly alkalic basalts: crystallization of spinel. Lithos 36,103^114.

Thy, P., Lofgren, G. E. & Imsland, P. (1991). Melting relations and theevolution of the Jan Mayen magma system. Journal of Petrology 32,303^332.

VanWagoner, N. A. & Leybourne, M. I. (1991). Evidence for magmamixing and a heterogeneous mantle on theWest Valley Segment ofthe Juan de Fuca Ridge. Journal of Geophysical Research. 96,16295^16318.

Walter, M. J. (1998). Melting of garnet peridotite and the origin of ko-matiite and depleted lithosphere. Journal of Petrology 39, 29^60.

Walter, M. J. & Presnall, D. C. (1994). Melting behavior of simplifiedlherzolite in the system CaO^MgO^Al2O3^SiO2^Na2O from 7 to35 kbar. Journal of Petrology 35, 329^359.

Walter, M. J., Sisson,T.W. & Presnall, D. C. (1995). A mass proportionmethod for calculating melting reactions and application to melt-ing of model upper mantle lherzolite. Earth and Planetary Science

Letters 135, 77^90.Wasylenki, L. E., Baker, M. B., Kent, A. J. R. & Stolper, E. M. (2003).

Near-solidus melting of the shallow upper mantle: partial meltingexperiments on depleted peridotite. Journal of Petrology 44,1163^1191.

Weaver, B., Kar, A., Davidson, J. & Colucci, M. (1996). Geochemicalcharacteristics of volcanic rocks from Ascension Island, SouthAtlantic Ocean. Geothermics 25, 449^470.

Weaver, B. L. (1991). The origin of ocean island basalt end-membercompositions: trace element and isotopic constraints. Earth and

Planetary Science Letters 104, 381^397.White, W. M. (1985). Sources of oceanic basalts: Radiogenic isotopic

evidence. Geology 13, 115^118.White, W. M. & Duncan, R. A. (1996). Geochemistry and geo-

chronology of the Society Islands: new evidence for deep mantlerecycling. In: Basu, A. & Hart, S. (eds) Earth Processes: Reading

the Isotopic Code, Am. Geophys. Union, Geophys. Monogr, . Washington:AGU 95, 183^206.

Willbold, M. & Stracke, A. (2006).Trace element composition of mantleend-members: Implications for recycling of oceanic and upper andlower continental crust. Geochemistry, Geophysics, Geosystems 7, doi:10.1029/2005GC001005.

Workman, R. K. & Hart, S. R. (2005). Major and trace element com-position of the depleted MORB mantle (DMM). Earth and

Planetary Science Letters 231, 53^72.Workman, R. K., Hart, S. R., Jackson, M., Regelous, M., Farley, K.

A., Blusztajn, J., Kurz, M. & Staudigel, H. (2004). Recycled meta-somatized lithosphere as the origin of the Enriched Mantle II(EM2) end-member: Evidence from the Samoan Volcanic Chain.Geochemistry, Geophysics, Geosystems 5, doi: 10.1029/2003GC000623.

Yasuda, A., Fujii,T. & Kurita, K. (1994). Melting phase relations of ananhydrous mid-ocean ridge basalt from 3 to 20GPa: Implicationsfor the behavior of subducted oceanic crust in the mantle. Journalof Geophysical Research 99, 9401^9414, doi: 10.1029/93JB03205.

Yaxley, G. M. & Green, D. H. (1998). Reactions between eclogite andperidotite: mantle refertilisation by subduction of oceanic crust.Schweizerische Mineralogische und Petrographische Mitteilungen 78,243^255.

Zindler, A. & Hart, S. (1986). Chemical Geodynamics. Annual Reviewof Earth and Planet Sciences 14, 493^571.

Zindler, A., Jagoutz, E. & Goldstein, S. (1982). Nd, Sr and Pb isotopicsystematics in a three-component mantle: a new perspective.Nature 298, 519^523.

APPENDIX : GEOCHEMICALMODEL INGFigure A1 is a flow chart in which the four main steps in themodeling process are identified.

PARAMETER IZAT ION FOR TRACEELEMENTSMelting and mixing equations are described using thefree parameters XDMM, XPeri, FRC, XOpx, XCpx, and XGrt,as well as two dependent parameters FOIB and XOL,as follows.


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Fig. A1. Flow chart describing steps 1^4 of the modeling procedure.


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Fig. A1. Continued.


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Melting of recycled crustThe composition of recycled crust-derived melt is ex-pressed using a non-modal accumulative fractional meltingequation (Shaw, 1970):

CiRCmelt ¼ Ci

RCFRC= 1� ð1� PiFRC=DiRCÞ

1=Pih i

ðA1Þ

where CiRCmelt and Ci

RC represent the abundances of elem-ent i in the recycled crust-derived melt and recycled crust,respectively, FRC is the degree of melting of recycled crust,and Di

RC and Pi are the bulk partition coefficients of elem-ent i in recycled crust and the phases entering the melt,respectively.Sub-solidus phase proportions in the recycled crust

(78·2% clinopyroxene, 17·8% garnet, 0·3% rutile, and3·7% quartz) were taken from Pertermann &Hirschmann (2003a) to calculate the bulk partition coeffi-cients in the recycled crust before melting.Residual phase abundances, which are essential to

obtain the melting reaction (Walter et al., 1995), were para-meterized in terms the degree of melting (FRC) basedon the volume fraction of each phase reported byPertermann & Hirschmann (2003a) as follows:

XGrt RC ¼ ð0 � 8605� FRCÞ=5 � 395 ðA2Þ

XCpx RC ¼ 1� FRC � XGrt RC ðA3Þ

where XGrt_RC and XCpx_RC are the fraction of garnet andclinopyroxene in the residual recycled crust, respectively.In general, rutile and silica minerals are present during

low-degree (520%) melting of MORB-like eclogite(Pertermann & Hirschmann, 2003b). At higher degrees ofmelting, in contrast, only clinopyroxene and/or garnet arethe residual phases at 2^3GPa. Therefore, for simplicitythe melting degree of the recycled crust was constrainedto be more than 20%. Partition coefficients between meltand silicate minerals were taken from Klemme et al.(2002, 2005) and are listed inTable 4.

Mixing of recycled crust-derived melt withthe model peridotiteThe hybrid source composition is expressed by mixing therecycled crust-derived melt and model peridotite:

CiSource ¼ Ci

RCmeltð1� XPeriÞ þ CiPeriXPeri ðA4Þ

where CiSource and Ci

Peri are the abundances of elementi in the OIB source and model peridodite respectively.XPeri is the weight fraction of the model peridotite compo-nent in the hybrid source (Fig. 14a).

Compositions of model peridotiteTo examine the effect of varying peridotite composition,the model peridotite compositions are expressed by

mixing two reference peridotite compositions, depletedMORB-source mantle (DMM, Workman & Hart, 2005)and primitive mantle (PM, Sun & McDonough, 1989)using the following equation:

CiPeri ¼ XDMMCi

DMM þ ð1� XDMMÞCiPM ðA5Þ

where CiPM, Ci

DMM, and XDMM are the abundances ofelement i in primitive mantle (PM), depleted MORBsource mantle (DMM), and the weight fraction of DMMin the model peridotite. Their trace element abundancesand present-day isotopic ratios are listed inTable 2.

Melting of hybrid source mantleMelting of the hybrid source is expressed using a modalbatch melting equation (Shaw, 1970). We treated FOIB, theweight fraction of melt produced by melting the hybridsource, as a dependent variable to reduce the number ofunknown parameters. In practice, the value of FOIB canbe determined to reproduce the Nb content of the St.Helena primary melt (33·6 ppm, Table 2) using the follow-ing equation based on modal batch melting:

FOIB ¼ CNbSource=ð33:6�DNb

R SourceÞ� �

=ð1�DNbR SourceÞ ðA6Þ

where DNbR Source is the bulk partition coefficient of Nb for

the residual hybrid source. Consequently, abundances ofelement i in the model OIB (Ci

OIB) can be expressedbased on the modal batch melting equation as follows:

CiOIB ¼ Ci

Source= FOIBð1�DiR SourceÞ þDi

R Source

� �ðA7Þ

where DiR Source is the bulk partition coefficient of element i

for the residual hybrid source. The bulk partition coeffi-cients are calculated using the mineral abundances in theresidual hybrid source and the partition coefficients listedinTable 4.With the exception of Pb, the partition coefficients for

the silicate phases used in equation (A7) were taken fromJohnson (1998), Salters et al. (2002), and Salters & Stracke(2004). As pointed out by several researchers (Salters et al.,2002; Hart & Gaetani, 2006), experimentally obtained par-tition coefficients for the major mantle silicate phases (seeTable 5) are lower for Pb than for Nd and Ce. This is incontradiction to the relatively constant Ce/Pb ratios (�25)of OIB, which imply similar bulk distribution coefficients(D) for Ce and Pb during melting. The most likely phasesto cause this discrepancy are sulfide minerals (McKenzie& O’Nions, 1998; Salters et al., 2002). Therefore, to high-light the effect of sulfide minerals, we set DPb to be theequal to DCe. If DPb is lower than DCe, the abundance ofPb in the recycled crust is higher than the results shownin Fig. 18 and Table 5. Therefore, the degree of Pb lossshown in Fig. 18 and Table 5 indicates minimum values.This has no effect on our conclusions that the recycled


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crust involved in the formation of St. Helena primary melthas lost Pb.

Mineral proportions in the residue frommelting of the hybrid source mantle:constrained from major elementsFour phases, olivine, orthopyroxene, clinopyroxene, andgarnet, were considered as potential residual phases duringthe melting of the hybrid source. To reliably estimate the re-sidual mineral proportions, we constrained the weight frac-tion of olivine in the residual hybrid source (XOL) withrespect to the other mineral phases, orthopyroxene (XOpx),clinopyroxene (XCpx), and garnet (XGrt), based on the massbalance of major element compositions.For the mass-balance calculations, we introduced the

component SiO2þ0·5AlO1·5 ^ 0·5NaO0·5 mol % (¼ SANmol %). The SAN mol % is a useful index to restrict themineral compositions. For example, olivine has an SAN of33mol %, whereas orthopyroxene, clinopyroxene andgarnet remain at 50mol %. Therefore, we can calculatethe mineral proportions without ambiguous mineralcompositions.By using the SAN mol %, we used the following

mass-balance relations to express two melting and onemixing event in the modeling scheme. The melting ofrecycled crust is given by

CSANRC ¼ FRCC

SANRCmelt þ CSAN

R RCð1� FRCÞ ðA8Þ

where CSANRC represents the SAN mol % of the recycled

crust, which is 52mol % assuming a MORB-like basalt(G2 basalt of Pertermann & Hirschmann, 2003a), andCSANR RC represents the SAN mol % of the residual recycled

crust. As described above, the residual minerals duringmelting of MORB-like eclogite are clinopyroxene� garnetat more than 20% melting. Because both clinopyroxeneand garnet have an SAN mol % of 50, the SAN of CSAN

R RCis 50mol % and equation (A8) is rearranged as

CSANRCmelt ¼ ð2þ 50FRCÞ=FRC: ðA9Þ

The mixing of recycled crust-derived melt and modelperidotite is expressed by

CSANSource ¼ XPeriC

SANPeri þ CSAN

RCmeltð1� XPeriÞ ðA10Þ

where CSANSource is the SAN of the hybrid source mantle.

Because both DMM and PM have an SAN of 40mol %,the model peridotite compositions assumed by mixingDMM and PM also have CSAN

Peri of 40mol %.In contrast, melting of the hybrid source is expressed as

CSANSource ¼ FOIBC

SANOIB þ CSAN

R Sourceð1� FOIBÞ: ðA11Þ

Olivine has an SAN of 33mol %, whereas the otherminerals, orthopyroxene, clinopyroxene, and garnet, havean SAN of 50mol %. In addition, the St. Helena primary

melt has CSANOIB ¼ 45·5mol %. Thus, the compositions of

the residual hybrid source are expressed using mineralabundances and their SAN values as follows:

CSANR Source ¼ 50ðXOpx þ XCpx þ XGrtÞ þ 33XOL: ðA12Þ

Therefore, by combining equations (A9)^(A12), the weightfraction of olivine in the residual hybrid source can be ex-pressed as

XOL ¼

40XPeri þ ð1� XPeriÞð2þ 50FRCÞ=

FRC � 45 � 5FOIB þ 50ðXOpx þ XCpx þ XGrtÞ

ðFOIB � 1Þ33ð1� FOIBÞ

:

ðA13Þ

Time evolution of radiogenic isotopic ratiosBy following the modeling scheme in Fig. 14, Sr, Nd, Pband Hf isotopic ratios are also calculated. The evolutionof the isotopic ratios of Bulk Earth, DMM, and oceaniccrust are calculated following the procedure of Strackeet al. (2003). A two-stage evolution model is applied to Sr,Nd and Hf isotopes, in which the initial isotopic compos-itions of the oceanic crust are assumed to be equal tothose of DMM at a certain age (¼ recycling age). In add-ition, the DMM was assumed to be isolated from the BulkEarth at a certain age (¼ the DMM age). The maximumDMM age is 3·4 Ga from Salters & Stracke (2004). Theminimum DMM age is set to 2 Ga, because the initial87Sr/ 86Sr of DMMat DMMage52 Ga becomes unreason-ably higher than the present-day 87Sr/ 86Sr of DMM.For Sr, Nd and Hf isotopic ratios, present-day Bulk

Earth values are set at 87Sr/ 86Sr¼ 0·7050,143Nd/144Nd¼ 0·512638, and 176Hf/177Hf¼ 0·28280 to-gether with 87Rb/ 86Sr¼ 0·09, 147Sm/144Nd¼ 0·1967, and176Lu/177Hf¼ 0·0335. The Sr, Nd and Hf isotopic ratios ofDMM are calculated using their present-day values(87Sr/ 86Sr¼ 0·70230, 143Nd/144¼ 0·51325, and176Hf/177Hf¼ 0·28326) and the given DMM age.In contrast, a single-stage evolution model was applied

to the evolution of Pb isotopic ratios (see Appendix B2 ofStracke et al., 2003). The Pb isotopic ratios of thepresent-day Bulk Earth are set at 206Pb/ 204Pb¼17·72,207Pb/ 204Pb¼15·49, and 208Pb/ 204Pb¼ 37·36. The Pb iso-topic ratios of DMM at a given DMM age are calculatedusing the Bulk Earth values at 4·55 Ga (206Pb/ 204Pb¼ 9·31,207Pb/ 204Pb¼10·3, 208Pb/ 204Pb¼ 29·48) and the parent/daughter (P/D) ratios of DMM (238U/ 204Pb¼ 8·2,232Th/ 238U¼ 3·8, and 232Th/ 204Pb¼ 31·2).Accordingly, the isotopic ratios of the recycled crust are

calculated based on P/D ratios in the recycled crust andthe initial isotopic ratio of the recycled crust. The isotopicratios in the model OIB are then calculated by mixing therecycled crust-derived melt with the peridotite. In the


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case of 87Sr/ 86Sr, for example, an isotopic mixture of modelOIB is calculated as follows:

87Sr86Sr

� �OIB¼

87Sr86Sr

� �Source¼ XPeri

CSrPeri

CSrSource

� � 87Sr86Sr

� �Peri

þð1� XPeriÞCSrRCmelt

CSrSource

� � 87Sr86Sr

� �RCmelt

:

ðA14Þ

For simplicity, we do not consider the isotopic evolutionthat would have occurred from the time the oceanic crustformed until it was subducted and became isolated withinthe mantle. In addition, we do not consider the enrichmentof Sr isotopic ratios in the altered oceanic crust. If we con-sider the alteration effect, the subducted oceanic crustshould have a lower Rb/Sr and the present P/D ratiosrequired by the recycled crust tend to decrease.


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evaluation of the oceanic crust-recycling model for himu oib

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