the origin and compositions of mesoarchean oceanic crust: evidence from the 3075 ma ivisaartoq...

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Lithos 100 (2008

The origin and compositions of Mesoarchean oceanic crust: Evidencefrom the 3075 Ma Ivisaartoq greenstone belt, SW Greenland

A. Polat a,⁎, R. Frei b,c, P.W.U. Appel d, Y. Dilek e, B. Fryer a,f,J.C. Ordóñez-Calderón a, Z. Yang f

a Department of Earth Sciences, University of Windsor, Windsor, ON, Canada N9B 3P4b Geological Institute, University of Copenhagen, 1350-Copenhagen, Denmarkc NordCEE, Nordic Center for Earth Evolution, Geological Institute, Denmarkd Geological Survey of Denmark and Greenland, 1350-Copenhagen, Denmark

e Department of Geology, Miami University, Oxford, OH 45056, USAf Great Lakes Institute for Environmental Research, University of Windsor, Windsor, ON, Canada N9B 3P4

Received 12 August 2006; accepted 8 June 2007Available online 2 August 2007

Abstract

The Mesoarchean (ca. 3075 Ma) Ivisaartoq greenstone belt contains well-preserved primary magmatic structures, such as pillowlavas, volcanic breccias, and clinopyroxene cumulate layers (picrites), despite the isoclinal folding and amphibolite faciesmetamorphism. The belt also includes variably deformed gabbroic to dioritic dykes and sills, actinolite schists, and serpentinites. TheIvisaartoq rocks underwent at least two stages of post-magmatic metamorphic alteration, including seafloor hydrothermal alteration andsyn- to post-tectonic calc-silicate metasomatism, between 3075 and 2961 Ma. These alteration processes resulted in the mobilization ofmany major and trace elements. The trace element characteristics of the least altered rocks are consistent with a supra-subduction zonegeodynamic setting and shallow mantle sources. On the basis of geological similarities between the Ivisaartoq greenstone belt andPhanerozoic forearc ophiolites, and intra-oceanic island arcs, we suggest that the Ivisaartoq greenstone belt represents a relic ofdismembered Mesoarchean supra-subduction zone oceanic crust. This crust might originally have been composed of a lower layer ofleucogabbros and anorthosites, and an upper layer of pillow lavas, picritic flows, gabbroic to dioritic dykes and sills, and dunitic towehrlitic sills.

The Sm–Nd and U–Pb isotope systems have been disturbed in strongly altered actinolite schists. In addition, the U–Pb isotopesystem in pillow basalts appears to have been partially open during seafloor hydrothermal alteration. Gabbros and diorites have the leastdisturbed Pb isotopic compositions. In contrast, the Sm–Nd isotope system appears to have remained relatively undisturbed in picrites,pillow lavas, gabbros, and diorites. As a group, picrites have more depleted initial Nd isotopic signatures (εNd=+4.23 to +4.97) thanpillow lavas, gabbros, and diorites (εNd=+0.30 to +3.04), consistent with a variably depleted, heterogeneous mantle source.

In some areas gabbros include up to 15 cm long white inclusions (xenoliths). These inclusions are composed primarily (N90%) ofCa-rich plagioclase and are interpreted as anorthositic cumulates brought to the surface by upwelling gabbroic magmas. Theanorthositic cumulates have significantly higher initial εNd (+4.8 to +6.0) values than the surrounding gabbroic matrix (+2.3 to +2.8),consistent with different mantle sources for the two rock types.© 2007 Elsevier B.V. All rights reserved.

Keywords: Archean; Greenstone belt; Oceanic crust; Pillow basalt; Anorthosite; Ocelli; Isotope

⁎ Corresponding author.E-mail address: [email protected] (A. Polat).

0024-4937/$ - see front matter © 2007 Elsevier B.V. All rights reserved.doi:10.1016/j.lithos.2007.06.021

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294 A. Polat et al. / Lithos 100 (2008) 293–321

1. Introduction

Archean greenstone belts are composed predominantlyof variably metamorphosed and deformed mafic to felsicvolcanic and siliciclastic sedimentary rocks (Condie, 1981;Goodwin, 1991; Eriksson et al., 1994; Condie, 2005).There are also volumetrically minor banded iron forma-tions (BIF), komatiites, gabbros, anorthosites, serpenti-nites, cherts, and carbonates (locally stromatolitic) in someArchean greenstone belts (Condie, 1981; Goodwin, 1991;Condie, 2005).Geochemical data derived from the study ofArchean greenstone belts over the last three decades showthe occurrence of diverse volcanic rock types, suggestingdiverse magmatic processes, such as plume and arcmagmatism, in oceanic or continental settings for theirorigin (Dostal andMueller, 1997; Polat et al., 1998; Kuskyand Polat, 1999; Polat and Hofmann, 2003; Dostal et al.,2004; Smithies et al., 2005a, b; Kerrich and Polat, 2006). Inaddition, greenstone belts from 3.8 to 2.5 Ga includevolcanic rock types reported from Phanerozoic convergentmargins, such as boninites, picrites, adakites, Mg-ande-

Fig. 1. (a)A simplified geologicalmap of the northeasternNuuk region, showing tbelt. Modified from Friend and Nutman (2005). (b) Geological map of the Ivisa

sites, and Nb-enriched basalts (Polat and Kerrich, 2006;and references therein). The distribution of rock types andthe internal structure of many Archean greenstone beltssuggest that they are the products of multiple geologicalprocesses, such as tectonism, magmatism, sedimentation,and metamorphism, operating over different spatial andtemporal scales (Corcoran and Dostal, 2000; Sandemanet al., 2004; Kerrich and Polat, 2006). Collectively, thegeological characteristics of many Archean greenstonebelts are comparable to those of lithotectonic assemblagesoccurring in Phanerozoic convergent plate boundaries(Kusky and Polat, 1999; Şengör and Natal’in, 2004;Kerrich and Polat, 2006; and references therein).

Thermal and geodynamic models and geochemicaland isotopic constraints derived from Archaean mafic–ultramafic rocks suggest that oceanic crust formationmust also have occurred in the Archean. From thermalmodeling, Abbott et al. (1994) inferred that Neoarcheanmid-ocean ridge crust was ∼11 km-thick, in contrast to∼7 km-thick in-situ oceanic crust developed at modernmid-ocean ridges. Numerical and modeling studies infer

heEoarchean toNeoarchean tectonic terranes and location of the Ivisaartoqartoq and surrounding area. Modified from Chadwick and Coe (1988).

Fig. 2. A simplified tectonostratigraphic column of the Ivisaartoq belt.Modified from Chadwick (1986, 1990).

295A. Polat et al. / Lithos 100 (2008) 293–321

that the negative buoyancy of ancient oceanic litho-sphere is responsible for b1% of its preservation in thePhanerozoic rock record, whereas relatively youngerand hotter backarc and forearc crusts have been morereadily accreted to orogenic belts (Cloos, 1993; VanHunen et al., 2002; Şengör and Natal’in, 2004).

The question of whether Archean greenstone beltsrepresent the fragments of ancient oceanic crust oralternatively are the remnants of continental flood basaltsremains controversial (Bickle et al., 1994; Kusky, 2004).Based on the criteria of xenocrystic zircons andgeochemical contamination trends of their mafic–ultramafic lavas, some greenstone belts may be consid-ered intra-oceanic in origin, whereas some others mayhave formed from mantle magmas erupted throughcontinental crust (Nisbet and Fowler, 1983; Polat et al.,1998; Arndt et al., 2001; Bleeker, 2002; Condie, 2005).

The recognition of Archean oceanic crust can be donemost effectively through comparative studies of the well-established lithological and geochemical characteristicsfor oceanic crustal fragments formed in Phanerozoicsupra-subduction zone environments. Phanerozoicsupra-subduction oceanic crust includes the presenceof: (1) contemporaneous mafic–ultramafic intrusive andextrusive units; (2) commonly picritic and boniniticunits; (3) chemical sedimentary rocks and sporadicallyvolcanogenic sediments; (4) diorite–plagiogranite intru-sives; (5) diagnostic metasomatic style temporallyassociated with the intrusions; (6) geochemical signa-tures of extrusive rock assemblages formed in intra-oceanic versus continental settings; (7) geochemical andisotopic signatures of a depleted mantle source; and (8)convergent margin geochemical signatures of magmas interms of REE/HFSE fractionations (see Dewey, 2003;Dilek, 2003; Hawkins, 2003; Schuth et al., 2004; Pythonet al., 2007).

Pillow structures, volcanic breccia, cumulate andocellar (eye-shaped) textures have been well preservedin low-strain domains of the Mesoarchean Ivisaartoqgreenstone belt in SW Greenland, despite the two majorphases of deformation and amphibolite-facies metamor-phism (Friend et al., 1981; Hall, 1981; Chadwick, 1985,1986; Hall et al., 1987; Chadwick, 1990; Appel, 1997;Polat et al., 2007). Preservation of these primarystructures and textures provides a unique opportunityto study Mesoarchean petrogenetic and geodynamicprocesses. The lithological, trace element, and hydro-thermal alteration characteristics of the Ivisaartoqgreenstone belt are comparable to those of Phanerozoicforearc ophiolites (Polat et al., 2007). Given the fact thatall rocks in the Ivisaartoq greenstone belt have beenmetamorphosed, the prefix ‘meta’ will be taken implicit.

In this study, we report new high-precision major andtrace element data (34 samples), and Nd (35 samples)and Pb (32 samples) isotope data obtained fromactinolite schists, ultramafic cumulates, gabbros, dior-ites, and anorthositic inclusions (xenoliths) in theIvisaartoq greenstone belt. Accordingly, the objectiveof this study is threefold: (1) to assess the effect of post-magmatic alteration on element mobility and of isotopiccomposition; (2) to understand the petrologic andgeodynamic origin of Mesoarchean oceanic crustpreserved in the Ivisaartoq greenstone belt; and (3) re-evaluate the existing geodynamic models proposed forthe origin of Archean anorthosites.

2. Regional geology and field characteristics

The Ivisaartoq greenstone belt contains the largestMesoarchean supra-crustal assemblage in southern WestGreenland (Fig. 1; Hall and Friend, 1979; Brewer et al.,1984; Chadwick, 1985, 1986, 1990; Friend and Nut-man, 2005). It is located in the central part of the innerGodthåbsfjord region (Fig. 1a). The belt occurswithin therecently recognized Mesoarchean (∼3075–2950 Ma)Kapisilik tectonic terrane (Friend and Nutman, 2005),which is tectonically bounded by the Eoarchean Isukasia


terrane (3600–3800 Ma) to the north, and the EoarcheanFæringehavn and the Neoarchean Tre Brødre terranes tothe south and west, respectively (Fig. 1a; Friend andNutman, 2005). The Kapisilik and Isukasia terranes werejuxtaposed and metamorphosed by 2950 Ma. It appears

Fig. 3. Field photographs of the Ivisaartoq pillow basalts, gabbros, cumulatesthe formation of stage I metasomatic assemblage during seafloor hydrothermwith plagioclase and quartz. (c) Pillow cores, rims and interstitial filled replacfine- and coarse-grained layers. (e) Clinopyroxene (cpx)-bearing cumulate. (

that the collision between the southern Færingehavn andthe Kapisilik terranes occurred at about 2800 Ma (Friendand Nutman, 2005). Field relationships indicate that theIsukasia terrane is structurally overlain by the Kapisilikterrane to the south; and the Kapisilik terrane is in turn

and serpentinites. (a) Compositionally zones pillow basalts, recordingal alteration. (b) Pillow basalts with concentric cooling fractures, filleded by stage I metasomatic assemblage. (d) Gabbro with well-preservedf) Deformed serpentinite in amphibolites (deformed pillow basalts).


structurally overlain by the Færingehavn and Tre Brødreterranes to the south–southwest.

The precise age of the volcanic and intrusive (gabbro,diorite) rocks in the Ivisaartoq greenstone belt is un-known. Siliceous volcaniclastic sedimentary rocks haveyielded an average U–Pb zircon age of 3075±15 Ma

ig. 4. Field photographs of the Ivisaartoq rocks. (a) Strongly deformed rock with a possible siliciclastic sedimentary origin. (b) Pyrite-bearingiliceous (metacherts) sedimentary layer within amphibolites. Contacts are typically sharp. (c) Tectonite with stage II calc-silicate assemblage near thewer and upper amphibolite contact. (d) Pillow basalts with drainage cavity-filling quartz and ocelli in the outer core. (e) Ocelli in an outer pillowore partly replaced by stage I calc-silicate assemblage. (f) Flattened centimeter-sized anorthositic inclusions in gabbros.

Fsloc

(Friend and Nutman, 2005; Polat et al., 2007), constrain-ing the maximum age of the belt. The Ivisaartoqgreenstone belt is intruded by weakly deformed 2961±12 Ma granites to the north, constraining the minimumage of the belt (Chadwick, 1990; Friend and Nutman,2005). The Ivisaartoq sequence is truncated by an up to


2 m-thick mylonite zone to the south, separating the beltfrom an association of leucogabbros and anorthosites.These leucogabbros and anorthosites are intruded by2963±8 Ma old tonalites and granodiorites (nowgneisses). On the basis of field observations and zirconages, Friend and Nutman (2005) interpreted the mylonite

Fig. 5. Field photographs of the Ivisaartoq rocks. (a) Flattened centimeter-sizealteration at a pillow basalt gabbro contact and in pillow cores. (c) Stage(d) Actinolite schist (dark) replaced by a massive layered stage II calc-silicate rvein. (f) Boudins of stage II calc-silicate assemblage in banded amphibolite.

zone as a post-2960 Ma structure deforming the Kapisilikterrane. The leucogabbro and anorthosite association islithologically and structurally similar to those found in theFiskenaesset region of southern West Greenland, and isinterpreted as intrusive into the Ivisaartoq greenstone belt(Chadwick, 1990).

d anorthositic inclusions (xenoliths) in gabbros. (b) Stage I calc-silicateII calc-silicate metasomatic assemblage, replacing actinolite schists.ock assemblage. (e) A diopside+garnet+hornblende+quartz ±epidote


The Ivisaartoq greenstone belt is composed mainly ofmafic to ultramafic volcanic rocks, gabbros, minor diorites,and serpentinites (Figs. 2–4; Hall, 1981; Chadwick, 1985,1986, 1990; Polat et al., 2007). Volcanic rocks consistdominantly of deformed pillow basalts and ultramafic lavaflows (Fig. 3; Chadwick, 1990; Polat et al., 2007). Sedi-mentary rocks constitute a volumetrically minor compo-nent of the belt (Figs. 1, 2, 4). Chadwick has subdivided theIvisaartoq greenstone belt into a lower and an upper am-phibolite unit (Fig. 2; Chadwick, 1985, 1986, 1990). Theseunits are separated by a thin layer (up to 50 m-thick) ofmagnetite-rich ultramafic schists, called the ‘magneticmarker’ (Fig. 2; Chadwick, 1986, 1990). Hydrothermalalteration of the ‘magneticmarker’ and volcanic rocks in itsvicinity resulted in the formation of calc-silicate rockshosting strata-bound scheelitemineralization (Appel, 1994,1997). The intensity of deformation appears to increasetowards the boundary between the two amphibolite units(Fig. 4c), suggesting that they are tectonically juxtaposed.

Volcanic rocks in the lower amphibolite unit are moreintensely deformed than those in their upper counterpart.They display a well-developed foliation characterized byamphibole- and plagioclase-rich domains. Pillow struc-tures are rare. Volcanic breccias are composed of pillowfragments locally with possible ocelli and hyaloclastites(Polat et al., 2007). Up to 50 m wide and 5 km long rustylayers of pyrite-rich siliceous rocks of probable felsicvolcaniclastic origin are exposed discontinuously in thelower amphibolite unit (Fig. 1b).

The upper amphibolite unit is composed mainly ofvariably deformed pillow basalts, actinolite schists, gab-bros, diorites, ultramafic cumulates, and serpentinites(Figs. 1–3). Serpentinites (ultramafic layers) are exposeddiscontinuously as three major layers throughout the se-quence (Figs. 1b, 2, 3f; Chadwick, 1986, 1990). Chadwick(1986) reports the presence of fresh olivine, likely of a

Table 1Mineralogical compositions of the Ivisaartoq rocks

Lithology Mineral assemblage

Cumulate Actinolite+clinopyroxene±plagioclase±qActinolite schist Actinolite+diopside±plagioclase±quartzInner pillow core Diopside+plagiocalse+quartz+epidote±Outer pillow core Hornblende+plagioclase+quartz±diopsidPillow rim Hornblende+quartz+plagioclase+epidoteGabbro Hornblende+plagioclase±epidote±quartzDiorite Hornblende+plagioclase±epidote±quartzAmphibolite Hornblende+plagioclase+quartz±diopsidInclusion in gabbros Plagioclase+hornblende±quartzOcelli in pillows Plagioclase+quartz+amphibole±epidoteCalc-silicate stage I Diopside+quartz+plagioclase+epidote±Calc-silicate stage II Diopside+garnet+amphibole+plagioclasMagnetic marker Actinolite+olivine+diopside+magnetite+

metamorphic origin, in serpentinites. On the basis of fieldrelationships, Chadwick (1986, 1990) suggested that theprotoliths of the serpentinites intruded as sills into sub-marine lavas. Inmany outcrops they are in tectonic contactwith pillow basalts and gabbros (Fig. 3f).

Pillow basalts are characterized bywell-preserved coreand rim structures (Fig. 3a). The least deformed pillowbasalts have concentric cooling cracks filled mainlywith quartz, and display way-up directions (Fig. 3b).Pillow cores are mineralogically zoned (Fig. 3a, c). Manyinner pillow cores display drainage cavities at the center,which are either empty or filled with quartz (Fig. 3c). Thepillow cores often display ocellar texture consistingchiefly of white ellipsoidal (eye-shaped) millimeter- tocentimeter-sized ocelli set in a dark green fine-grainedmafic matrix (Fig. 4d, e). The contacts between ocelli andmatrix are sharp. In many cores the ocelli-matrix texturehas been partly to completely replaced by a calc-silicatemetasomatic assemblage (Fig. 4). Pillow rims oftendisplay silica alteration; some pillows have beencompletely silicified. Some pillows are composed pre-dominantly of actinolite, consistent with an ultramaficprotolith. Primary magmatic textures, such as clinopyr-oxene cumulates, are locally preserved in ultramafic flowsof low-strain domains (Fig. 3e). With increasing intensityof deformation clinopyroxene cumulates grade intoactinolite schists.

Gabbros and minor diorites occur as one to severaltens of meter-thick sills and dykes in pillow basalts(Fig. 3d). Gabbroic and dioritic sills also occursporadically between pillow basalts and ultramaficflows. Chilled margins between pillow basalts andgabbroic dykes are preserved in a few locations. Primaryigneous textures and minerals are locally preserved in low-strain domains (Fig. 3). Some gabbros contain deformedanorthositic inclusions up to 15 cm long (Figs. 4f, 5a). Like

uartz

amphibole±sulphide± titanitee±epidote± titanite±biotite± titanite

±biotitee±epidote± titanite±sulphide

hornblende±scapolitee+quartz±vesuvianite±scapolite±epidote± titanite±calcite±scheeliteplagioclase±epidote±scapolite±calcite± titanite±scheelite


pillow basalts, gabbros and diorites underwent calc-silicatemetasomatic alteration mainly along fractures and pillowbasalt contacts (Fig. 5b). An alternating association of over200 m-thick actinolite schist, pillow basalt, gabbro, andpyrite-bearing siliceous volcaniclastic rocks is exposed inthe northernmargin of the belt. This sequence is intensively

Fig. 6. Photomicrographs of the Ivisaartoq rocks. (a) Altered clinopyroxene(b) Actinolite schist (crossed polarized light). (c) Gabbro (crossed polarized(e) Magnetic marker (plain polarized light). Magnetite occurs mainly alongdiopside and calcite (plain polarized light). (Cpx: clinopyroxene; act: actino

sheared and metasomatized mainly along the actinolitelayers (Fig. 5c). The intensity of calc-silicate metasomaticalteration increases from gabbros through pillow basalts toactinolite schists (Fig. 5). Siliceous pyrite-bearing rockswere interpreted as metamorphosed cherts (Chadwick,1990; Fig. 4b). There are minor, up to several meter-thick,

phenocrysts surrounded by actinolite matrix (crossed polarized light).light). (d) Ocelli texture in outer pillow core (plain polarized light).the foliation planes. (f) Calc-silicate assemblage consisting mainly oflite; hornb: hornblende; plag: plagioclase).

Fig. 7. Photomicrographs of anorthositic inclusions in gabbros (seeFig. 5a), showing recrystallized plagioclase.

Table 2Measured and recommended values for the USGS standards BHVO-1and BHVO-2

Element BHVO-2(n=15)

BHVO-2 BHVO-1(n=10)

BHVO-1

Measured Recommended Measured Recommended

Li 4.5 5.0 4.63V 331 317 398.1Cr 286 280 340.3Co 45 45 44.55Ni 116 119 148.69Cu 135 127 143.11Zn 102 103 170.43Rb 9.0 9.8 9.12Sr 390 389 390.5Y 24.1 26.0 23.2 28Zr 165 172 163.8 179Nb 15.29 18.00 15.3 19Cs 0.10 0.10 0.13Ba 131 130 127.1 139La 14.93 15 14.7 16Ce 37.24 38 37.2 39Pr 5.27 5.17 5.4Nd 24.08 25.0 23.6 25Sm 5.95 6.20 5.99 6.4Eu 2.00 1.97 2.06Gd 6.17 6.30 6.21 6.4Tb 0.92 0.90 0.90 0.96Dy 5.19 5.13 5.2Ho 0.95 1.04 0.94 0.99Er 2.54 2.49Tm 0.33 0.32 0.33Yb 1.94 2.00 1.91 2.00Lu 0.27 0.28 0.27 0.29Hf 4.39 4.10 3.91 4.4Ta 1.01 1.40 0.88 1.2Pb 1.77 1.85 2.6Th 1.47 1.20 1.36 1.1U 0.36 0.27


lenses of siliciclastic sedimentary rocks in the upper unit(Fig. 4a). Contacts between volcanic and sedimentaryrocks are sharp (Fig. 4b).

Field and petrographic observations suggest that theIvisaartoq greenstone belt underwent at least two stages ofcalc-silicate metasomatic alteration prior to the intrusionof 2961 Ma granitoids (Polat et al., 2007). Stage Ialteration assemblage typically occurs within the innerpillow cores, pillow interstitials, and along the pillow-gabbro contacts (Figs. 3–5; Table 1). This assemblage isalso found along the pillow-gabbro contacts and fractureswithin gabbros (Fig. 5b). Stage IImetasomatic assemblageoccurs as calc-silicate veins and boudins that areconcordant to discordant to the dominant foliation planesin the replaced host rocks, consistent with a syn- to postdeformation origin (Fig. 5). Most of these veins arespatially associated with shear zones.

3. Petrography

Themineralogical characteristics of different rock typesare summarized in Table 1 and photomicrographs of major

mineral assemblages are shown in Fig. 6. Cumulates arecomposed primarily of altered clinopyroxene phenocrysts(Figs. 4e, 6a; Table 1). Ultramafic schists are composeddominantly of actinolite (Fig. 6b; Table 1). Gabbros anddiorites are composedmainly of hornblende+plagioclase±epidote±quartz (Fig. 6c).

Ocelli in the pillow cores (Figs. 4d, e; 6d) consistmainly of plagioclase (30–50%)+quartz (30–40%)+amphibole (10–20%)±epidote (0–5%). No internalstructure has been observed in the ocelli. The darkermatrix surrounding the ocelli is made of amphibole (50–60%)+plagioclase (20–30%)+quartz (10–20%)±epi-dote (0–5%)± titanite (0–5%) (Fig. 6d; Table 1). Thepillow rims are composed of fine-grained hornblende+quartz+plagioclase±epidote±biotite.

Magnetic marker is composed primarily of actino-lite + olivine (metamorphic) +magnetite + diopside ±


garnet±plagioclase±cheelite± titanite (Fig. 6e). StageI metasomatic assemblage is composed predominantlyof epidote (now mostly diopside)+quartz+plagioclase±hornblende±scapolite (Table 1). Stage II metasomaticassemblage consists mainly of diopside+garnet+amphi-bole+plagioclase+quartz±vesuvianite±scapolite±epi-dote± titanite± calcite± scheelite (Table 1; Fig. 6f).Amphibolites (foliated pillows) are composed mainly ofhornblende+plagioclase+quartz±diopside±epidote±titanite±sulphide.

White inclusions in gabbros have an assemblageof plagioclase (90%)+amphibole (5–10%)+quartz (0–5%), consistent with an anorthositic composition(Fig. 7; Table 1). Because of extensive recrystallization,magmatic plagioclase is rarely present. Amphiboles inthe anorthositic inclusions typically have dark green toblue green pleochroism and range from subhedral toeuhedral.

Table 3Summary of major (wt.%) and trace (ppm) element concentrations and sign

Actinolite schist Cumulates Pillow lavas Gabb

SiO2 (wt.%) 44.2–55.4 48.7–50.3 47.7–55.7 47.6–TiO2 0.10–0.67 0.27–0.28 0.37–0.75 0.50–Al2O3 4.0–15.2 6.3–7.1 8.7–15.2 13.6–Fe2O3 6.8–12.4 9.2–10.4 7.7–13.9 9.0–MgO 15.5–25.6 22.3–23.5 4.5–18.8 7.8–CaO 4.5–12.4 9.6–10.0 8.5–17.2 8.3–Mg-number 72.7–86.7 81.7–83.2 53.6–76.9 53.9–Cr (ppm) 1325–12700 1575–1670 62–5700 230–Co 68–107 80–84 48–96 45–Ni 430–1520 730–830 125–705 87–Sc 12–50 24–26 26.5–43.4 33–V 100–550 100–170 136–485 190–Nb 0.16–2.68 0.69–0.77 0.09–1.61 0.10–Zr 11–25 12.5–15.6 22.3–42.4 28.2–Th 0.06–0.30 0.22–0.36 0.37–0.79 0.22–Y 2.6–21.6 7.0–8.4 9.4–15.2 13.6–La 0.23–80.0 1.59–1.83 2.25–3.56 2.21–Nd 0.63–54.0 2.35–2.85 3.66–5.26 4.12–Sm 0.26–7.42 0.69–0.87 1.10–1.90 1.26–Gd 0.42–6.62 0.98–1.22 1.52–2.47 1.83–Yb 0.34–2.44 0.80–1.00 1.08–1.69 1.49–La/Ybcn 0.47–23.5 1.27–1.60 1.09–2.19 0.80–La/Smcn 0.64–7.74 1.50–1.90 0.97–2.31 0.90–Gd/Ybcn 0.84–1.91 0.97–1.02 1.00–1.20 1.00–Eu/Eu⁎ 0.77–2.16 0.61–0.81 0.63–1.05 0.70–Al2O3/TiO2 23–38 23.5–25.6 20.3–26.3 15.8–Nb/Ta 4.7–17.4 11.5–16.9 13.3–16.3 8.7–Y/Ho 3.9 –72.0 25.5–28.5 25.7–28.7 24.5–Zr/Y 1.1–6.0 1.8–2.1 2.0–2.9 1.2–Ti/Zr 59–163 103–128 76–116 94–Zr/Zr⁎ 0.1–1.9 0.66–0.78 0.52–0.94 0.50–Nb/Nb⁎ 0.01–0.96 0.31–0.41 0.28–0.63 0.30–Ti/Ti⁎ 0.25–1.78 0.64–0.81 0.64–0.82 0.60–

4. Analytical methods and data presentation

All whole-rock samples were powdered using an agatemill in the Department of Earth Sciences, University ofWindsor, Canada. Major and some trace elements (Zr, Sc,Ni) were determined by Thermo Jarrel-Ash ENVIRO IIICP at ACTLABS in Ancaster, Canada. Samples weremixed with a flux of lithium metaborate and lithiumtetraborate, and were fused in an induction furnace.Molten sample was immediately poured into a solution of5% nitric acid containing an internal standard, and wasmixed continuously until completely dissolved. Totals ofmajor element oxides are 100±1 wt.% and the analyticalprecisions are 1 to 2%.

Samples were analyzed for REE, HFSE, LILE, andtransition metals (Co, Cr, and V) by a high-sensitivityThermo Elemental X7 ICP-MS in the Great LakesInstitute for Environmental Research (GLIER),University

ificant element ratios for the Ivisaartoq rocks

ros Diorites Anorthositic inclusions Gabbroic matrix

51.0 55.2–57.1 47.4–49.0 46.9–50.41.00 0.64–1.14 0.04–0.08 0.73–1.0816.0 15.0–16.9 29.1–30.3 15.8–16.413.2 6.2–8.0 2.4–3.2 10.7–13.014.1 3.8–7.6 1.1–1.6 8.1–8.612.3 9.2–12.0 14.0–15.9 10.6–11.175.6 50.6–70.1 43.6–49.5 56.6–60.61060 180–1030 6–11 210–24061 36–48 6–9 45–50230 60–120 b2 90–14041 35–48 1–3 38–42475 230–300 14–21 240–2841.77 1.33–3.06 0.10–0.21 1.45–2.1548.5 36.3–68.4 3.6–21.1 48–570.74 0.15–0.36 0.04–0.12 0.33–0.3819.8 9.5–22.2 1.3–2.2 18–207.98 2.44–6.63 0.68–0.98 2.67–3.099.47 4.50–10.30 0.61–1.00 6.38–7.252.47 1.34–2.76 0.15–0.24 2.04–2.362.94 1.70–3.47 0.19–0.32 2.90–3.212.20 1.08–2.58 0.16–0.21 2.07–2.293.30 1.05–2.05 2.80–3.32 0.92–1.002.30 0.89–1.89 2.60–3.35 0.92–0.971.40 1.10–1.30 0.94–1.28 1.12–1.161.00 0.89–1.07 4.1–6.5 0.84–0.9427.0 14.8–23.5 390–673 15–2216.7 12.6–16.0 12.7–14.627.0 23.5–27.1 29–32 24–262.7 1.2–3.8 1.8–9.8 2.6–2.9123 95–106 13–90 17–230.90 0.86–1.00 0.50–3.05 0.86–0.950.80 0.30–0.90 0.11–0.16 0.53–0.820.93 0.80–1.00 0.40–0.64 0.69–0.91

Table 4Major (wt.%) and trace (ppm) element concentrations and significant element ratios for the Ivisaartoq actinolite schists, mafic to ultramafic pillows,cumulates, gabbros, and diorites

Actinolite schists

485426 485427 485430 485431 485433 485434 455435 485436 485437

SiO2 (wt.%) 52.35 52.67 46.39 53.35 44.17 45.67 46.47 45.38 50.84TiO2 0.21 0.21 0.44 0.10 0.30 0.39 0.31 0.67 0.47Al2O3 5.67 5.59 12.68 3.98 7.40 11.48 9.46 15.22 10.92Fe2O3 9.11 8.96 11.52 6.84 11.99 12.41 12.28 9.63 9.28MnO 0.13 0.14 0.22 0.12 0.18 0.15 0.20 0.23 0.16MgO 21.83 21.28 15.52 22.59 25.58 19.22 21.18 24.09 15.53CaO 10.12 10.54 11.44 12.42 10.02 9.57 9.52 4.51 10.36K2O 0.02 0.01 0.30 0.04 0.01 0.15 0.05 0.01 0.04Na2O 0.54 0.59 1.46 0.56 0.33 0.93 0.50 0.24 2.39P2O5 0.01 0.01 0.03 0.01 0.02 0.03 0.02 0.02 0.02LOI 3.11 2.76 1.78 2.76 3.27 0.39 0.90 0.75 2.04Mg-number 82.6 82.5 72.7 86.7 80.9 75.4 77.4 83.2 76.8Cr (ppm) 12733 12484 1988 6715 11438 7298 11896 1325 4794Co 86 84 68 84 99 102 94 107 87Ni 1249 1192 477 1520 1241 1106 767 431 877Rb 0.2 0.2 11.8 0.2 0.3 8.7 0.5 1.8 1.2Sr 29 30 28 48 190 32 89 30 226Cs 0.1 0.1 3.1 0.1 0.3 2.6 0.1 1.0 0.3Ba 5 5 56 35 13 35 17 10 38Sc 20.6 18.5 25.4 12.3 25.6 31.2 32.4 49.9 35.7V 199 207 287 100 260 269 338 550 414Nb 0.38 0.41 1.11 0.16 0.50 0.72 2.68 0.77 0.78Ta 0.02 0.03 0.08 0.01 0.04 0.05 0.57 0.05 0.04Zr 12.4 11.3 28.1 10.7 11.3 19.6 20.5 24.7 22.3Th 0.16 0.12 0.06 0.07 0.20 0.06 0.30 0.09 0.09U 0.26 0.44 0.02 0.03 0.25 0.08 0.37 0.08 0.08Y 4.8 4.9 2.6 7.5 6.8 8.1 21.6 3.7 3.7La 2.179 2.842 3.440 0.229 0.687 22.238 2.659 80.022 0.912Ce 4.132 6.216 7.244 0.860 2.044 44.179 6.741 151.484 2.220Pr 0.447 0.693 0.972 0.108 0.332 4.915 0.928 16.151 0.357Nd 1.802 2.538 4.464 0.634 1.858 17.358 3.967 53.958 1.802Sm 0.440 0.522 1.332 0.255 0.702 2.630 1.029 7.415 0.572Eu 0.201 0.268 0.482 0.083 0.245 0.945 0.543 3.602 0.439Gd 0.641 0.684 1.671 0.424 1.037 1.683 1.241 5.644 0.669Tb 0.127 0.137 0.304 0.071 0.193 0.252 0.211 0.777 0.128Dy 0.846 0.859 1.995 0.502 1.393 1.384 1.562 4.515 0.795Ho 0.194 0.196 0.410 0.103 0.295 0.277 0.326 0.945 0.159Er 0.613 0.604 1.317 0.337 0.952 0.843 0.994 2.795 0.515Tm 0.092 0.095 0.177 0.048 0.135 0.121 0.152 0.391 0.069Yb 0.630 0.576 1.132 0.347 0.964 0.816 1.059 2.443 0.419Lu 0.100 0.098 0.174 0.052 0.152 0.143 0.167 0.357 0.068Cu 38.1 36.7 3.0 19.0 44.9 2.64 5.06 2.72 2.4Zn 70.1 59.6 139.0 170.6 112.1 99.78 181.90 143.58 78.0Ga 14.1 13.7 34.1 15.9 17.8 25.6 23.7 28.1 25.7Pb 1.0 0.9 1.9 5.2 6.4 5.6 2.6 2.3 6.0La/Ybcn 2.48 3.54 2.18 0.47 0.51 19.53 1.80 23.49 1.56La/Smcn 3.55 3.90 1.85 0.64 0.70 6.06 1.85 7.74 1.14Gd/Ybcn 0.84 0.98 1.22 1.01 0.89 1.70 0.97 1.91 1.32Ce/Ce⁎ 1.03 1.09 0.97 1.34 1.05 1.04 1.05 1.03 0.95Eu/Eu⁎ 1.16 1.37 0.99 0.77 0.88 1.28 1.47 1.64 2.16Al2O3/TiO2 26.6 27.3 28.8 37.9 24.8 29.2 30.3 22.6 23.4Nb/Ta 15.2 13.6 13.6 11.4 12.4 15.9 4.7 16.4 17.4Y/Ho 24.8 24.7 6.4 72.0 23.0 29.2 66.1 3.9 23.3Zr/Y 2.6 2.3 2.7 4.1 1.5 2.9 2.5 1.1 6.0Ti/Zr 103 108 94 59 158 120 91 163 125

(continued on next page)



Actinolite schists

485426 485427 485430 485431 485433 485434 455435 485436 485437

Zr/Zr⁎ 0.96 0.68 0.80 1.85 0.69 0.20 0.70 0.09 1.52Nb/Nb⁎ 0.12 0.12 0.25 0.96 0.81 0.02 0.95 0.01 0.77Ti/Ti⁎ 0.95 0.81 0.70 0.76 0.83 0.44 0.65 0.25 1.78North 64°44.903′ 64°44.903′ 64°44.900′ 64°44.900′ 64°44.940′ 64°44.940′ 64°44.940′ 64°44.943′ 64°44.962′West 049°53.261′ 049°53.261′ 049°53.466′ 049°53.466′ 049°53.402′ 049°53.402′ 049°53.402′ 049°53.379′ 049°53.313′

Cumulates Pillow lavas

485473 a 485474 a 485475 a 485478 485418 a 485420 a 485422 485424 485468 485469 485470

SiO2 50.34 48.73 48.84 48.41 52.06 47.72 55.661 50.09 50.95 51.74TiO2 0.27 0.28 0.27 0.43 0.51 0.45 0.516 0.60 0.56 0.54Al2O3 6.36 7.05 6.28 10.87 12.01 10.58 12.489 14.12 14.49 14.32Fe2O3 9.21 10.42 9.90 12.51 11.17 12.31 7.724 11.60 10.85 10.33MnO 0.17 0.17 0.17 0.22 0.19 0.20 0.227 0.19 0.17 0.17MgO 23.09 23.51 22.29 14.90 10.12 17.27 4.513 9.75 8.90 8.59CaO 10.23 9.56 12.03 11.26 12.66 9.64 17.204 10.78 11.27 11.94K2O 0.02 0.01 0.01 0.15 0.01 0.90 0.050 0.14 0.14 0.16Na2O 0.30 0.25 0.20 1.23 1.23 0.89 1.585 2.67 2.63 2.16P2O5 0.01 0.02 0.01 0.02 0.04 0.04 0.030 0.05 0.05 0.05LOI 4.15 4.64 3.97 1.50 0.82 2.93 0.776 0.85 0.69 0.60Mg-number 83.2 81.7 81.7 70.2 64.2 73.5 53.6 62.5 61.9 62.2Cr 1575 1670 1654 1617 5713 2999 5282 2325 230 300 300Co 84 84 83 80 93 62 83 53 57 55 57Ni 845 852 735 857 605 331 647 235 125 128 141Rb 0.8 0.7 0.7 0.6 12.3 0.4 148.5 8.7 2.0 2.0 2.7Sr 33 31 33 36 43 90 16 69 123 120 96Cs 0.2 0.2 0.1 0.2 4.7 0.1 29.4 0.9 0.0 0.0 0.0Ba 1.9 10.9 6.1 2.45 45 30 137 23 37 58 49Sc 24 24 26 35.5 37.3 35.0 38.3 37.3 37.3 36.2V 104 112 107 172 178 199 173 421 234 294 291Nb 0.77 0.73 0.69 0.67 0.94 1.26 1.09 1.32 1.54 1.51 1.42Ta 0.05 0.05 0.06 0.05 0.07 0.09 0.07 0.09 0.11 0.09 0.10Zr 12.5 14.6 15.6 22.3 33.3 27.5 35.3 36.3 30.2 28.2Th 0.34 0.26 0.36 0.23 0.40 0.708 0.55 0.70 0.79 0.48 0.48U 0.293 0.668 0.183 0.271 0.101 0.192 0.167 0.194 0.190 0.183 0.145Y 7.0 7.1 8.4 7.04 11.0 12.3 10.9 13.4 14.7 14.4 14.1La 1.69 1.59 1.82 1.83 2.25 3.31 2.39 3.53 2.98 2.85 2.84Ce 4.10 3.75 4.26 3.95 4.95 7.83 5.60 8.00 7.60 7.08 6.99Pr 0.52 0.54 0.63 0.53 0.76 1.06 0.77 1.10 1.09 1.01 1.00Nd 2.35 2.43 2.89 2.44 3.66 4.96 3.69 5.14 5.26 4.94 4.83Sm 0.73 0.70 0.87 0.69 1.11 1.39 1.11 1.48 1.67 1.56 1.43Eu 0.23 0.22 0.21 0.22 0.44 0.46 0.38 0.53 0.58 0.57 0.56Gd 0.98 0.98 1.22 1.01 1.64 1.85 1.59 1.92 2.17 1.99 1.90Tb 0.18 0.17 0.22 0.17 0.29 0.33 0.28 0.35 0.39 0.36 0.37Dy 1.20 1.26 1.52 1.19 1.96 2.28 1.89 2.31 2.65 2.56 2.44Ho 0.26 0.27 0.33 0.25 0.42 0.48 0.41 0.50 0.57 0.55 0.51Er 0.85 0.82 0.99 0.79 1.30 1.46 1.27 1.53 1.73 1.59 1.61Tm 0.13 0.12 0.14 0.12 0.19 0.21 0.18 0.22 0.23 0.23 0.23Yb 0.84 0.81 1.03 0.82 1.24 1.40 1.19 1.51 1.50 1.47 1.49Lu 0.13 0.12 0.15 0.12 0.18 0.21 0.18 0.23 0.22 0.22 0.21Cu 10.8 14.4 6.7 18.3 15.5 78.3 8.0 38.5 43.1 40.7 94.9Zn 79.2 91.2 80.6 81.7 93.5 108.7 98.9 68.0 104.7 99.1 91.2Ga 15.0 17.8 16.1 16.2 30.7 32.2 46.0 32.3 39.6 42.2 40.3Pb 1.6 1.7 1.8 0.33 2.2 1.7 1.3 2.0 5.3 3.2 11.2La/Ybcn 1.45 1.41 1.27 1.60 1.30 1.70 1.44 1.68 1.42 1.39 1.37La/Smcn 1.65 1.62 1.50 1.89 1.45 1.70 1.55 1.71 1.28 1.32 1.42Gd/Ybcn 0.97 1.00 0.98 1.02 1.09 1.09 1.11 1.05 1.19 1.12 1.06Ce/Ce⁎ 1.07 1.00 0.97 0.98 0.93 1.02 1.01 0.99 1.03 1.02 1.02

Table 4 (continued )


Eu/Eu⁎ 0.82 0.81 0.61 0.78 0.99 0.88 0.88 0.97 0.93 0.98 1.05Al2O3/TiO2 23.6 25.1 23.5 25.1 23.5 23.7 24.2 23.5 25.9 26.3Nb/Ta 16.9 13.9 11.5 14.7 14.4 14.1 15.7 14.5 14.2 16.3 14.6Y/Ho 27.5 26.5 25.5 28.5 25.9 25.6 26.4 26.6 25.7 26.2 27.4Zr/Y 1.8 2.1 1.9 2.0 2.7 2.5 2.6 2.5 2.1 2.0Ti/Zr 128 115 103 116 92 97 88 99 110 116Zr/Zr⁎ 0.66 0.78 0.68 0.77 0.88 0.94 0.88 0.85 0.75 0.74Nb/Nb⁎ 0.37 0.41 0.31 0.29 0.36 0.30 0.40 0.32 0.49 0.49 0.46Ti/Ti⁎ 0.78 0.83 0.64 0.79 0.78 0.79 0.72 0.75 0.75 0.78North 64°

44.763′64°44.763′

64°44.763′

64°44.761′

64°.375′

64°44.334′

64°44.361′

64°44.325′

64°44.798′

64°44.812′

64°44.812′

West 049°51.318′

049°51.318′

049°51.318′

049°51.332′

049°.130′

049°.183′

049°56.678′

049°.342′

049°51.544′

049°.477′

049°51.477′

Pillow lavas Gabbros Diorites

485481 a 485482 a 485486 485423 485477 499730A 485432 499739 485429

SiO2 50.39 49.75 48.31 50.98 47.55 57.02 55.18TiO2 0.35 0.37 0.37 0.62 1.00 0.65 0.64Al2O3 8.69 9.02 8.70 16.04 15.83 14.96 15.16Fe2O3 11.08 10.65 13.91 9.76 12.34 6.22 8.02MnO 0.20 0.18 0.25 0.16 0.21 0.12 0.13MgO 18.61 17.10 18.77 8.08 9.21 7.58 6.99CaO 8.46 11.50 8.62 11.87 11.50 10.14 11.40K2O 1.33 0.03 0.08 0.04 0.53 0.23 0.11Na2O 0.86 1.38 0.96 2.40 1.76 3.01 2.31P2O5 0.03 0.02 0.03 0.04 0.07 0.05 0.05LOI 2.05 1.80 2.46 0.72 1.02 0.81 1.06Mg-number 76.9 76.1 72.8 62.1 0.6 0.71 63.3Cr 1298 1414 1640 3153 323 232 842 407 1029Co 81 81 81 75 53 57 45 36 42Ni 608 659 705 378 127 152 87 60 93Rb 275.8 2.3 1.0 4.5 1.7 63.5 1.5 2.9 1.6Sr 69 76 61 46 54 90 51 62.8 107Cs 60.8 0.5 0.6 0.7 0.1 14.1 0.1 0.1 0.1Ba 216 10 31 44 12 61 29 66.1 59Sc 27.6 26.5 27.7 37 41 n.d. 47.4 48V 142 136 204 485 213 276 475 235 247Nb 1.11 1.29 1.20 1.40 1.68 2.13 1.86 1.33 1.56Ta 0.08 0.09 0.09 0.09 0.10 0.24 0.13 0.11 0.11Zr 25.5 29.5 27.7 22.4 33.2 48.5 36.29 40.4Th 0.68 0.83 0.72 0.66 0.50 0.34 0.52 0.79 0.69U 0.148 0.286 0.260 0.228 0.26 0.09 0.19 0.18 0.15Y 10.9 10.7 9.4 14.1 13.94 19.55 15.38 9.45 11.64La 3.56 3.34 3.31 3.56 3.40 2.93 2.44 3.09 4.32Ce 7.80 7.96 7.52 8.86 8.43 8.40 7.68 7.28 10.04Pr 1.00 1.04 0.93 1.19 1.22 1.34 1.22 1.00 1.31Nd 4.34 4.57 4.02 5.53 5.57 6.71 5.99 4.52 6.04Sm 1.10 1.17 1.19 1.60 1.56 2.14 1.96 1.34 1.74Eu 0.41 0.41 0.28 0.60 0.60 0.76 0.68 0.53 0.59Gd 1.54 1.54 1.52 2.15 2.22 2.96 2.61 1.70 2.14Tb 0.25 0.27 0.25 0.37 0.36 0.53 0.46 0.29 0.37Dy 1.84 1.75 1.70 2.44 2.42 3.56 3.06 1.96 2.39Ho 0.38 0.39 0.36 0.55 0.52 0.76 0.63 0.40 0.49Er 1.19 1.15 1.06 1.67 1.53 2.31 1.86 1.16 1.40Tm 0.18 0.18 0.16 0.24 0.22 0.34 0.27 0.17 0.18Yb 1.24 1.21 1.08 1.54 1.49 2.20 1.66 1.08 1.33



Cumulates Pillow lavas

485473 a 485474 a 485475 a 485478 485418 a 485420 a 485422 485424 485468 485469 485470


Lu 0.18 0.19 0.16 0.25 0.23 0.33 0.27 0.16 0.19Cu 2.6 5.2 25.7 84.5 44.4 86.5 19.0 22.8 263.2Zn 92.3 78.3 110.7 97.2 116.4 77.5 69.2 54.5 66.8Ga 52.6 23.9 24.9 36.7 38 31 37 27.6 45Pb 1.9 1.3 2.8 1.7 7.46 1.41 4.44 4.9 2.54La/Ybcn 2.06 1.98 2.19 1.66 1.64 0.95 1.05 2.05 2.33La/Smcn 2.31 2.05 1.99 1.59 1.56 0.96 0.89 1.63 1.79Gd/Ybcn 1.03 1.05 1.16 1.16 1.23 1.11 1.30 1.30 1.33Ce/Ce⁎ 1.01 1.05 1.05 1.06 1.02 1.04 1.09 1.02 1.03Eu/Eu⁎ 0.96 0.93 0.63 0.99 0.99 0.91 0.92 1.07 0.94Al2O3/TiO2 25.0 24.1 23.8 25.9 15.8 23.1 23.5Nb/Ta 13.3 14.2 13.5 15.2 16.7 8.7 14.6 12.7 14.4Y/Ho 28.7 27.6 25.9 25.4 27.0 25.8 24.5 23.74 23.5Zr/Y 2.3 2.8 2.9 2.4 2.5 1.2 3.8 3.5Ti/Zr 81 76 79 112 123 106 95Zr/Zr⁎ 0.81 0.88 0.88 0.52 0.78 0.87 1.01 0.86Nb/Nb⁎ 0.26 0.28 0.31 0.37 0.46 0.78 0.89 0.38 0.31Ti/Ti⁎ 0.66 0.68 0.64 0.79 0.93 1.00 0.79North 64°44.515′ 64°.515′ 64°.438′ 64°.355′ 64°44.767′ 64 44.492 64°44.742 64°44.776′West 049°53.299′ 049°53.361′ 049°54.399′ 049°56.753′ 049°51.576′ 049°56.613 049°54.082 049°53.759′a Published in Polat et al. (2007).


Pillow lavas Gabbros Diorites

485481 a 485482 a 485486 485423 485477 499730A 485432 499739 485429


of Windsor, Canada. Wet chemical procedures wereconducted under clean lab conditions, and all acids weredistilled twice. Approximately 100–130 mg of samplepowder was used for dissolution. Samples were dissolvedin a concentrated HF–HNO3 mixture at a temperature of∼120 °C for four days, and then further attacked withconcentrated HNO3 until no residue was visible. BHVO-1and BHVO-2 were used as international referencematerials to estimate precision and accuracy (Table 2).Analytical precisions are estimated as follows: 1 to 10 %for REE, Rb, Li, Cs, Sr, Ba, Y, Nb, Cu, Zn, and Pb; 10 to20% for Zr, V, Cr, Co, andU; and 20 to 30% for Ta and Th(Table 2).

Selected elements are normalized to primitive mantle(pm) (Hofmann, 1988) and chondrite (cn) (Sun and Mc-Donough, 1989). Nb/Nb⁎, Zr/Zr⁎, Ce/Ce⁎, and Eu/Eu⁎

ratios, representing anomalies, were calculated with respectto the neighboring immobile elements, following themethod of Taylor and McLennan (1985). Samples wererecalculated to 100 % anhydrous for inter-comparisons.Mg-numbers (%) were calculated as the molecular ratio ofMg/(Mg+Fe2+), where Fe2+ is assumed to be 90% of totalFe.

Whole-rock Pb and Sm–Nd isotope analyses werecarried out on a VG Sector 54-IT TIMS in the GeologicalInstitute, University of Copenhagen, Denmark. Dissolu-tion of the powder samples was achieved in two suc-cessive, but identical steps, which consisted of a strong 8NHBr attack that has been shown to effectively dissolve

accessory phosphates (Frei et al., 1997; Schaller et al.,1997), followed by a concentrated HF–14N HNO3 mix-ture and finally by strong 9N HCl. Independent dissolu-tions were performed for REE and Pb analyses. A mixed150Nd–149Sm spike was added to the REE aliquot before-hand. Chemical separation of REEs was carried out onconventional cation exchange columns, followed by anSm–Nd separation using HDEHP-coated beads (BIO-RAD) charged in 6 ml quartz glass columns. Neodymiumratios were normalized to 146Nd/144Nd=0.7219. The meanvalue for our internal JM Nd standard (referenced againstLa Jolla) during the period of measurement was 0.511098for 143Nd/144Nd, with a 2σ external reproducibility of ±0.000011 (seven measurements). Procedural blanks runduring the period of these analyses show insignificant blanklevels of ∼5 pg Sm and ∼12 pg Nd. Precisions forconcentration analysis are approximately 0.5% for Sm andNd. Initial εNd values were calculated at 3075 Ma U–Pbzircon ages obtained from volcanoclastic rocks (Friend andNutman, 2005).

5. Geochemical results

5.1. Major and trace elements

5.1.1. Clinopyroxene cumulates (picrites) and actinoliteschists

Field relationships suggest that actinolite schists werederived from clinopyroxene cumulates by increasing


intensity of deformation and metamorphism. Cumulateshave more uniform major and trace element composi-tions than actinolite schists (Tables 3, 4; Figs. 8, 9).Cumulates have sub-chondritic Nb/Ta (11.5–16.9) andZr/Y (1.8–2.1) ratios, and slightly super-chondriticAl2O3/TiO2 (23–25) ratios (Tables 3, 4). On chondrite-and primitive mantle-normalized diagrams, they have thefollowing trace element characteristics: (1) moderately

Fig. 8. Chondrite-normalized REE and primitive mantle-normalized traceChondrite normalization values are from Sun and McDonough (1989) and p

enriched LREE (La/Smcn=1.50–1.90; La/Ybcn=1.30–1.60) patterns; (2) flat HREE (Gd/Ybcn=0.97–1.02)patterns; and (3) negative Eu (Eu/Eu⁎=0.61–0.82), Nb(Nb/Nb⁎=0.29–0.41), Zr (Zr/Zr⁎=0.66–0.78), and Ti (Ti/Ti⁎=0.64–0.83) anomalies (Fig. 8a, e; Table 4).

Despite their simple mineralogical composition(Fig. 6b; Table 1), actinolite schists display largevariations in Al2O3 (4.0–15.2 wt.%), Cr (1325–

element patterns for cumulates, pillow lavas, gabbros, and diorites.rimitive mantle normalization values are from Hofmann (1988).


12500 ppm), Ni (430–1250 ppm), V (100–550 ppm),TiO2 (0.10–0.67 wt.%), Nb (0.16–2.70 ppm), and Y(2.6–21.6 ppm) (Tables 3, 4). They have moderatevariations in SiO2 (44–53 wt.%), MgO (15.5–25.6 wt.%),Fe2O3 (6.9–12.3wt.%), Zr (11–25 ppm), Ga (14–34 ppm),and Co (84–106 ppm). Abundances of REE (e.g.La=0.23–80 ppm; Ce=0.9–152 ppm) are extremelyscattered (Table 4). On a chondrite-normalized diagram,they have the following characteristics: (1) variablydepleted to strongly enriched LREE patterns (La/Smcn=0.6–7.7; La/Ybcn=0.50–23.50); (2) slightly deplet-ed to enriched HREE (Gd/Ybcn=0.84–1.90) patterns; (3)negative to positive Eu (Eu/Eu⁎=0.9–2.16) anomalies; and(4) positive Ce (Ce/Ce⁎=0.97–1.34) anomalies (Fig. 9).On a primitive mantle-normalized diagram, they havethe following significant features: (1) variably negativeNb (Nb/Nb⁎=0.02–0.96); and (2) negative to positiveTi (Ti/Ti⁎=0.2–1.8) and Zr (Zr/Zr⁎=0.09–1.85) anoma-lies (Fig. 9). Al2O3/TiO2 (22–38) ratios are chondritic tosuper-chondritic, whereas the ratios of Zr/Y (1.4–4.0) andTi/Zr (59–162) range from sub-chondritic to super-chondritic values (Tables 3, 4).

5.1.2. Pillow lavas, gabbros, and dioritesPillow lavas are basaltic, but have a variable composi-

tion (Tables 3, 4). Ti/Zr (76–116) and Zr/Y (2.03–2.94)ratios range fromsub-chondritic to slightly super-chondriticvalues. Al2O3/TiO2 (24–25) ratios are slightly super-chon-dritic. The ratios of Nb/Ta (13.3–16.3) and Y/Ho (25.6–28.7) tend to be sub-chondritic. In addition, they have thefollowing trace element characteristics: (1) flat to enrichedLREE (La/Smcn=0.97–2.31; La/Ybcn=1.10–2.20) pat-terns; (2) flat to slightly enriched HREE (Gd/Ybcn=1.03–1.19) patterns; and (3) negative Nb (Nb/Nb⁎=0.26–0.63),

Fig. 9. Chondrite-normalized REE and primitive mantle-normalized trace elemantle-normalized diagram reflect the mobility of LREE. Given that sampleNb, Zr, and Ti concentrations, we suggest that these elements were less mMcDonough (1989) and primitive mantle normalization values are from Ho

Zr (Zr/Zr⁎=0.52–0.94), and Ti (Ti/Ti⁎=0.64–0.82)anomalies (Fig. 8).

The following compositional ranges in gabbros anddiorites represent the new and previously published(Polat et al., 2007) data (Table 3, 4). Major elementcompositions of these rocks are similar to those of pillowlavas (Tables 3, 4). There are large variations in Ni (121–234 ppm), Cr (246–1060 ppm), and REE (e.g., La=2.2–8.0 ppm), andmoderate variations in Co (52–61 ppm), V(186–264), Zr (28–44 pm), and Y (14–20 ppm) (Tables3, 4). Mg-numbers vary between 54 and 76 (Tables 3, 4).The ratios of Al2O3/TiO2 (16–27), Zr/Y (2.1–2.7), andTi/Zr (94–121) extend from sub-chondritic to super-chondritic values (see Sun and McDonough, 1989). Nb/Ta (8.7–16.7) and Y/Ho (23.5–27.0) ratios are sub-chondritic. In addition, they have the following traceelement characteristics: (1) slightly depleted to moder-ately enriched LREE (La/Smcn = 0.90–2.30; La/Ybcn=0.80–3.20) patterns; (2) flat to slightly enrichedHREE (Gd/Ybcn=1.00–1.37) patterns; and (3) variablylarge negative Nb (Nb/Nb⁎=0.29–0.90) and Ti (Ti/Ti⁎=0.60–0.90) anomalies (Fig. 8; Table 3, 4). Dioriteshave more evolved geochemical compositions (i.e.higher SiO2, but lower MgO, Fe2O3, Ni andCr concentrations and Mg-numbers) than gabbros(Tables 3, 4). The chondrite- and primitive mantle-nor-malized trace element patterns of diorites are similar tothose of gabbros (Fig. 8).

5.1.3. Anorthositic inclusions and surrounding gabbroicmatrix

The anorthositic inclusions have high concentra-tions of Al2O3 (29–30 wt.%) and CaO (14–16 wt.%)(Table 5). They have extremely low Ni (b2 ppm),

ment patterns for actinolite schists. Non-coherent patterns in primitivewith very different REE patterns (and concentrations) have similar Th,obile than REEs. Chondrite normalization values are from Sun andfmann (1988).

Table 5Major (wt.%) and trace (ppm) compositions and significant element ratios of anorthositic inclusions and surrounding gabbroic matrix

Anorthositic inclusions Gabbroic matrix

499728-A1 499729-A1 499729-B1 499731-A1 499731-B1 499731-C1 499728-A2 499729-A2 499729-B2 499731-A2 499731-B2 499731-C2

SiO2 47.39 48.00 48.99 48.46 48.84 48.44 50.19 49.85 48.30 49.88 46.89 50.41TiO2 0.04 0.05 0.05 0.05 0.05 0.08 0.91 0.73 0.98 0.94 1.08 0.99Al2O3 30.09 30.36 29.43 30.16 29.10 29.52 15.81 15.81 16.44 15.98 16.40 15.79Fe2O3 3.16 2.42 3.06 2.93 3.24 3.17 11.10 11.77 11.95 11.46 13.02 10.66MnO 0.05 0.04 0.04 0.06 0.05 0.05 0.19 0.19 0.20 0.19 0.21 0.19MgO 1.47 1.20 1.35 1.14 1.56 1.38 8.43 8.19 8.56 8.12 8.56 8.27CaO 15.21 15.87 13.95 14.23 14.09 14.68 10.55 11.11 10.75 10.89 10.87 11.18K2O 0.16 0.03 0.17 0.16 0.18 0.06 0.55 0.26 0.66 0.22 0.78 0.30Na2O 2.42 2.02 2.94 2.80 2.89 2.62 2.19 2.05 2.09 2.23 2.10 2.14P2O5 0.01 0.02 0.02 0.01 0.01 0.01 0.08 0.04 0.07 0.07 0.09 0.07LOI 0.44 0.49 0.36 0.37 0.31 0.01 0.98 0.80 1.09 0.63 1.12 1.33Mg-number 48.0 49.5 46.7 43.6 48.7 46.3 60.1 58.0 58.7 58.4 56.6 60.6Cr (ppm) 6.2 5.7 8.3 5.6 5.8 10.9 209 222 236 240 227 231Co 9.4 8.0 7.4 6.4 8.8 7.8 45 47 49 46 50 45Ni b2 b2 b2 b2 b2 b2 121 131 142 141 101 91Rb 8.4 1.9 18.4 16.8 15.1 3.5 74 18 100 9 112 19Sr 393 399 436 429 408 383 133 134 140 142 128 142Cs 1.88 0.42 4.28 4.57 3.71 1.28 15.2 3.1 21.5 1.2 24.0 3.3Ba 62 25 64 54 78 44 207 38 253 44 292 80Sc 1.00 2.01 2.01 b1 2.01 3.00 39 38 42 40 42 41V 15.3 17.1 17.4 16.4 14.0 20.9 244 265 275 260 284 265Ta n.d. n.d. n.d. n.d. n.d. n.d. 0.13 0.11 0.16 0.15 0.16 0.16Nb 0.192 0.178 0.123 0.124 0.103 0.208 1.87 1.45 2.12 1.89 2.15 2.04Zr 5.0 3.6 3.6 5.0 4.2 7.5 47.5 52.4 56.6 50.3 52.6 52.7Th 0.054 0.038 0.036 0.128 0.042 0.038 0.37 0.38 0.38 0.35 0.43 0.33U 0.017 0.017 0.028 0.022 0.031 0.020 0.11 0.06 0.08 0.07 0.07 0.07Y 2.21 2.03 1.46 1.32 1.72 2.03 17.9 18.9 19.8 19.1 20.0 19.3La 0.950 0.986 0.682 0.720 0.779 0.946 2.67 2.94 3.09 2.85 3.08 2.96Ce 1.962 1.718 1.274 1.291 1.727 1.735 8.39 8.56 8.96 8.66 9.47 8.75Pr 0.225 0.245 0.166 0.157 0.190 0.213 1.25 1.29 1.41 1.32 1.44 1.36Nd 0.959 1.005 0.698 0.619 0.776 0.931 6.38 6.56 7.14 6.71 7.25 6.84Sm 0.235 0.253 0.187 0.153 0.196 0.241 2.04 2.15 2.29 2.17 2.36 2.25Eu 0.392 0.363 0.393 0.367 0.355 0.365 0.68 0.77 0.73 0.73 0.76 0.73Gd 0.288 0.281 0.203 0.190 0.242 0.322 2.81 2.90 3.13 3.01 3.21 3.03Tb 0.052 0.054 0.033 0.029 0.038 0.049 0.50 0.51 0.55 0.53 0.56 0.54Dy 0.321 0.313 0.223 0.186 0.246 0.344 3.30 3.52 3.67 3.53 3.75 3.61Ho 0.071 0.067 0.046 0.042 0.054 0.071 0.73 0.74 0.77 0.77 0.77 0.77Er 0.207 0.219 0.154 0.146 0.164 0.227 2.15 2.22 2.38 2.25 2.38 2.33Tm 0.031 0.031 0.025 0.022 0.025 0.032 0.30 0.32 0.34 0.33 0.34 0.34Yb 0.205 0.212 0.175 0.167 0.173 0.208 2.07 2.08 2.25 2.18 2.29 2.23


309A.Polat

etal.

/Lithos

100(2008)

293–321

Anorthositic inclusions Gabbroic matrix

499728-A1 499729-A1 499729-B1 499731-A1 499731-B1 499731-C1 499728-A2 499729-A2 499729-B2 499731-A2 499731-B2 499731-C2

Lu 0.032 0.030 0.026 0.025 0.026 0.031 0.31 0.31 0.34 0.32 0.35 0.33Cu 30.9 16.8 34.1 18.2 37.1 26.9 26 27 23 31 25 54Zn 36.8 23.7 23.7 49.8 18.0 23.5 76 79 79 103 112 118Ga 38.6 35.5 37.0 34.8 37.7 33.5 43 27 51 27 55 30Pb 4.7 4.1 3.5 5.2 3.4 3.4 2.00 2.34 2.72 3.97 3.52 4.99Li 27.0 18.5 30.2 29.4 31.6 24.9 77 63 89 59 99 64La/Ybcn 3.32 3.33 2.80 3.09 3.23 3.26 0.92 1.01 0.99 0.94 0.96 0.95La/Smcn 2.87 2.77 2.60 3.35 2.83 2.81 0.92 0.97 0.95 0.93 0.92 0.92Gd/Ybcn 1.16 1.10 0.96 0.94 1.16 1.28 1.12 1.15 1.15 1.14 1.16 1.12Eu/Eu⁎ 4.58 4.12 6.11 6.54 4.96 4.00 0.86 0.93 0.82 0.87 0.84 0.84Ce/Ce⁎ 1.04 0.86 0.93 0.94 1.10 0.95 1.12 1.08 1.05 1.09 1.10 1.07Al2O3/TiO2 674 563 561 653 592 391 17.3 21.7 16.8 17.0 15.2 16.0Zr/Y 9.58 1.76 2.49 7.98 2.45 3.67 2.65 2.77 2.86 2.64 2.63 2.73Ti/Zr 13 91 86 26 70 60 116 84 104 110 121 111Ti/V 17 19 18 17 21 22 23 17 21 21 22 22Nb/Nb⁎ 0.16 0.12 0.13 0.11 0.11 0.15 0.82 0.53 0.74 0.75 0.80 0.76Zr/Zr⁎ 3.05 0.48 0.69 2.36 0.74 1.09 0.89 0.95 0.95 0.90 0.86 0.91Ti/Ti⁎ 0.40 0.48 0.63 0.64 0.53 0.64 0.90 0.69 0.86 0.85 0.91 0.88North 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′ 64° 44.492′West 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′ 49° 56.613′

n.d.: not determined.


310A.Polat

etal.

/Lithos

100(2008)

293–321


Cr (5.6–10.9 ppm), Co (6.4–9.4 ppm), Sc (1–3 ppm), V (14–21 ppm), and TiO2 (0.04–0.08 wt.%) contents. Mg-numbers range between 44 and50. In addition, they display very low HFSE (Nb=0.10–0.21 ppm; Y=1.3–2.2 ppm) and REE (La=0.68–0.98 ppm; Yb= 0.17–0.22 ppm) concen-trations (Tables 3, 5). Al2O3/TiO2 (390–670) ratiosare extremely high. On primitive mantle- and chon-drite-normalized diagrams, they have the followingsignificant features: (1) moderately enriched LREE(La/Smcn=2.60–3.35; La/Ybcn=2.8–3.3); (2) flat toslightly enriched HREE (Gd/Ybcn=0.94–1.28); (3)large positive Eu (Eu/Eu⁎=4.1–6.5) anomalies; and(4) negative Nb (Nb/Nb⁎=0.11–0.16) and Ti (Ti/Ti⁎=0.4–0.6) anomalies (Fig. 10).

The gabbroic matrix has higher MgO, Fe2O3, TiO2,Sc, Ni, Cr, and Co, but lower CaO, Al2O3, and Srcontents than the anorthositic inclusions (Tables 3, 5).In addition, the matrix is characterized by higherabsolute concentrations of REE and HFSE than theinclusions (Tables 3, 5). In comparison to theinclusions, the matrix has less fractionated LREE(La/Smcn=0.92–0.97 versus 2.60–3.35) patterns, andsmaller Nb (Nb/Nb⁎=0.53–0.82 versus 0.12–0.16),and Ti (Ti/Ti⁎=0.69–0.91 versus 0.40–0.64) anoma-lies (Fig. 10; Tables 3, 5).

Fig. 10. Chondrite-normalized REE and primitive mantle-normalizedtrace element patterns for inclusions (xenoliths) in gabbros (seeFig. 5a). Chondrite normalization values are from Sun andMcDonough (1989) and primitive mantle normalization values arefrom Hofmann (1988).

5.2. Nd isotopes


Regression of the Sm–Nd isotope data for clinopyr-oxene cumulates and their more deformed counterpartactinolite schists yields an errorchron age of 3092±260 Ma (MSWD=97) (Fig. 11a; Table 6; excludingextremely altered actinolite schist samples 485434 and485436). This age, within uncertainties, is in goodagreement with the 3075±15 Ma U–Pb zircon age ofthe spatially associated siliceous volcaniclastic sedi-mentary rocks (see Friend and Nutman, 2005). Largeuncertainty in the errorchron age is likely due to largescatter in the data. Cumulates have a narrow range ofinitial εNd (+4.97 to +4.23) values, whereas actinoliteschists display large variations (+2.54 to +9.53)(Table 6). Samples (e.g., 485434, 485436) with verylarge εNd (+8.35 and +9.53) values have much higherLREE concentrations (Tables 4, 6; Nd=16–43 ppm)and more fractionated LREE (La/Ybcn=20–24; La/Smcn=6.1–7.7) patterns, compared to the rest of thesamples in the group (Fig. 9; Table 4). In addition, thesesamples display the lowest 147Sm/144Nd (0.088–0.089)ratios in the group (Table 6).

5.2.2. Pillow lavas, gabbros, and dioritesPillow lavas, gabbros, and diorites define an

errorchron age of 3069 ± 220 Ma (MSWD=80)(Fig. 11b). This age, within uncertainties, agrees wellwith the 3075±15 Ma U–Pb zircon age of siliceousvolcaniclastic sedimentary rocks (Friend and Nutman,2005; Polat et al., 2006). Large uncertainty in theerrorchron ages reflects the narrow compositionalrange and large scatter in the data points. The initialεNd values in pillow lavas (+1.10 to +3.10) overlapwith, but extend to higher values than, gabbros anddiorites (+0.30 to +2.39) (Table 6). Samples withhigher MgO content tend to have greater initial εNdvalues (e.g., 485475, 485477). All rock types havesimilar range of 147Sm/144Nd (Table 6).

5.2.3. Anorthositic inclusions and surrounding gab-broic matrix

The Sm–Nd isotopic compositions of the anorthositicinclusions and surrounding gabbroic matrix are pre-sented in Table 7. The inclusions have much higherinitial εNd values than the matrix (εNd=+4.8 to +6.0versus +2.3 to +2.8). The Nd isotopic compositions ofthe inclusions are comparable to those of clinopyroxenecumulates (Tables 6, 7). The initial εNd (+2.3 to +2.8)isotopic composition of the matrix overlaps with, but

Fig. 11. 147Sm/144Nd verses 143Nd/144Nd and 206Pb/204Pb versus 207Pb/204Pb errorchron diagrams for cumulates, actinolite schists, pillow lavas,gabbros, and diorites. The isoplot program of Ludwig (1988, 2003) was used for age and initial 143Nd/144Nd ratio calculations.


extends to higher values than, gabbros (devoid ofanorthositic inclusions) and diorites (+0.3 to +2.4).

5.3. Pb isotopes


On a 207Pb/204Pb versus 206Pb/204Pb isotope diagram,cumulates and actinolite schists define an errorchron ageof 2774±180 (MSWD=167) (Fig. 11c). This age is lowerthan the U–Pb zircon age (3075 Ma) of the spatiallyassociated volcanoclastic sedimentary rocks, but agrees,within uncertainties, with 2781 and 2847MaU–Pb zirconages yielded by granodioritic gneisses to the west of theIvisaartoq greenstone belt (Friend and Nutman, 2005).Cumulates have higher and more limited rages of206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb than actinoliteschists (Table 8).

5.3.2. Pillow lavas, gabbros, and dioritesPillow lavas, gabbros, and diorites define an errorchron

age of 3058±240 Ma (MSWD=505) (Fig. 11d). Despite

the large errors, this age agrees with the Sm–Nd (3069±220Ma) errorchron and U–Pb zircon (3075±15Ma) ages.Pillow lavas display larger variations in 206Pb/204Pb and207Pb/204Pb ratios than gabbros and diorites (Table 8).Gabbros tend to have larger 206Pb/204Pb and 207Pb/204Pbratios than diorites, consistent with higherU/Pb ratios in thelatter group (Table 8).

6. Discussion

6.1. Post-magmatic alteration, element mobility andmodification of isotopic composition

The Ivisaartoq greenstone belt underwent at least twostages of metamorphic alteration prior to the intrusion of2961±12 Ma granitoids, resulting in the formation ofwidespread calc-silicate metasomatic mineral assemblages(Figs. 3–6; Polat et al., 2007). Stage I metasomatic assem-blage appears to have formed during seafloor hydrother-mal alteration under greenschist to lower-amphibolitefacies metamorphic conditions (Polat et al., 2007). StageII metasomatic assemblage was formed during a regional

Table 6Sm–Nd isotope composition of the Ivisaartoq actinolite schists, cumulates, pillow lavas, gabbros, and diorites

Sample # Rock type 143Nd/144Nd ±2σ Nd (ppm) 147Sm/144Nd Sm (ppm) εNd (3075 Ma) Sm/Nd

485426 Actinolite schist 0.512027 11 1.82 0.1557 0.469 4.37 0.26485430 Actinolite schist 0.512244 9 3.45 0.1709 0.974 2.54 0.28485431 Actinolite schist 0.513339 28 0.72 0.2236 0.267 3.06 0.37485433 Actinolite schist 0.511848 11 3.67 0.1489 0.902 3.56 0.25485434 Actinolite schist 0.510883 13 16.3 0.0894 2.413 8.35 0.15485435 Actinolite schist 0.511767 11 3.59 0.1465 0.869 2.91 0.24485436 Actinolite schist 0.510923 9 42.6 0.0884 6.218 9.53 0.15485437 Actinolite schist 0.512678 11 1.53 0.1869 0.471 4.71 0.31485473 a Cpx cumulate 0.512504 10 2.07 0.1795 0.615 4.23 0.30485474 a Cpx cumulate 0.512512 11 2.44 0.1787 0.719 4.72 0.30485475 a Cpx cumulate 0.512479 12 2.77 0.1764 0.807 4.97 0.29485481 a Pillow basalt 0.512086 11 4.56 0.1618 1.219 3.10 0.27485482 a Pillow basalt 0.511967 10 4.34 0.1584 1.135 2.10 0.26485486 Pillow basalt 0.512060 13 3.95 0.1646 1.075 1.44 0.27485414 Pillow basalt 0.512617 14 5.67 0.1906 1.786 2.03 0.31485418 a Pillow basalt 0.512549 10 3.13 0.1853 0.958 2.81 0.31485420 a Pillow basalt 0.512226 12 4.59 0.1732 1.313 1.29 0.29485422 Pillow basalt 0.512438 9 3.88 0.1792 1.149 3.04 0.30485468 Pillow basalt 0.512453 12 5.19 0.1848 1.586 1.10 0.31485469 Pillow basalt 0.512483 12 4.86 0.1850 1.486 1.65 0.31485428 Gabbro 0.512922 9 6.0 0.204 2.025 2.86 0.34485432 Gabbro 0.512472 10 5.72 0.1878 1.776 0.30 0.31485438 Gabbro 0.511830 12 8.28 0.1521 2.080 1.91 0.25485467 Gabbro 0.512601 13 3.97 0.1925 1.263 0.94 0.32485472 Gabbro 0.512449 6 4.99 0.1854 1.528 0.79 0.31485476 Gabbro 0.512395 11 5.49 0.1813 1.643 1.39 0.30485477 Gabbro 0.512248 11 5.49 0.1727 1.565 1.94 0.29499730-A Gabbro 0.512693 5 6.94 0.1935 2.218 2.39 0.32499739 Diorite 0.512219 5 4.64 0.1735 1.332 1.01 0.29485483 Diorite 0.512098 10 9.97 0.1663 2.739 1.55 0.27485429 Diorite 0.512158 10 5.57 0.1702 1.566 1.15 0.28485484 Diorite 0.512188 11 6.84 0.1713 1.937 1.28 0.28

All initial εNd ages calculated at 3075 Ma yielded by U–Pb zircon analyses (see Friend and Nutman, 2005).a Published in Polat et al. (2007).


tectonothermal metamorphic event undermiddle- to upper-amphibolite facies metamorphic conditions (Appel, 1997;Polat et al., 2007).

Many pillow basalts aremineralogically and chemicallyzoned (Fig. 3a; Polat et al., 2007). The rims have highercontents of Fe2O3, MgO, MnO, and K2O, whereas theinner and outer cores possess higher concentrations ofCaO, and Na2O and SiO2, respectively, consistent with themobility of these elements during post-magmatic alter-ation. Similarly, large variations inBa, Sr, Pb,Rb,Cs, Li,U,Zn, and Cu contents between pillow cores and rims areconsistent with the mobility of these elements. Comparedwith the less altered cores, the rims have lower LREEabundances and La/Smcn ratios, indicating the loss of theseelements. In contrast to the above elements, Al2O3, TiO2,Th, Zr, Y, Cr, Ni, Co, Ga, and HREE display minor varia-tions between the cores and rims, suggesting that theseelements were relatively immobile during Mesoarcheanseafloor hydrothermal alteration. Similarly, REE, HFSE

(Ti, Nb, Ta, Zr, Y) in gabbros, diorites, pillow lavas, andcumulates display coherent primitive mantle- and chon-drite-normalized patterns (Fig. 8), indicating that theseelements were also relatively immobile during post-mag-matic alteration.

Cumulates with relict clinopyroxene phenocrysts arecharacterized bymore coherent trace element patterns andnarrower ranges of many major and trace elements thantheir more deformed actinolite schist counterparts(Figs. 8, 9). In addition, they have a narrow range ofSm/Nd (0.29–0.30) ratios and initial εNd (+4.23 to +4.97)values (Table 6), consistent with the Sm–Nd system re-maining closed.

Given the preservation of primary minerals andtexture in cumulates, the initial εNd values in these rockslikely reflect the near-primary magmatic composition(Fig. 3e; Table 6).

In contrast to those in cumulates, many elements inactinolite schists display large variations (Fig. 9; Table 4).

Table 7Sm–Nd isotope composition of the anorthositic inclusions and surrounding gabbroic matrix in the Ivisaartoq belt

Sample # Rock type 143Nd/144Nd ±2σ Nd (ppm) 147Sm/144Nd Sm (ppm) εNd (3075 Ma)

499728-A1 Anorthositic inclusion 0.511941 6 0.850 0.1505 0.2114 4.77499729-A1 Anorthositic inclusion 0.511935 7 0.980 0.1474 0.2388 5.85499731-A1 Anorthositic inclusion 0.511802 7 0.578 0.1406 0.1343 5.99499731-B1 Anorthositic inclusion 0.511811 6 0.615 0.1420 0.1444 5.57499731-C1 Anorthositic inclusion 0.512015 6 0.871 0.1532 0.2204 5.14499728-A2 Gabbroic matrix 0.512699 6 6.732 0.1929 2.146 2.69499729-A2 Gabbroic matrix 0.512687 6 6.612 0.1929 2.107 2.49499731-A2 Gabbroic matrix 0.512732 5 6.936 0.1943 2.227 2.79499731-B2 Gabbroic matrix 0.512713 4 7.269 0.1934 2.323 2.81499731-C2 Gabbroic matrix 0.512703 4 7.223 0.1942 2.317 2.30

All initial εNd ages calculated at 3075 Ma yielded by U–Pb zircon analyses (see Friend and Nutman, 2005).


Despite the fact that actinolite schist samples analyzed forthis study came from the least metasomatized outcrops,they still display a large spread in U, Pb, Sm, Nd con-centrations, and isotopic ratios (Tables 4, 6, 8). Addition-ally, these samples have large positive to negative Eu andCe anomalies (Tables 3, 4). Even the samples that werecollected from the same outcrop (e.g., 485426 and485427; and 485430 and 485431; 485434 and 485435)display significant variations in these isotopic ratios(Tables 4, 6, 8). In summary, these geochemical charac-teristics are consistent with the mobility of U, Pb, Sm, andNd in actinolite schists during post-magmatic alteration.

The late Archean (2774±180) Pb–Pb errorchron age(Fig. 11c) yielded by actinolite schists may reflect the∼2800 Ma tectonothermal event that affected the region(Friend and Nutman, 2005). Despite the Mesoarchean(3092±260Ma) Sm–Nd errorchron age (Fig. 11a), somesamples (e.g., 485434 and 485435) collected from thesame outcrop have significant variations in Sm/Nd ratios(0.15–0.24) and initial εNd (+2.9 to +8.35), indicatingthat the Sm–Nd isotope system in these rocks waspartially open on a whole-rock scale during Mesoarch-ean hydrothermal alteration. Compared to their lessstrained cumulate counterparts, actinolite schists appearto have been significantly enriched in LREE (e.g., La/Ybcn=16–24) and LILE (e.g., Rb=12 ppm) elements byhydrothermal fluids. Accordingly, we interpret the largevariations in the initial εNd (+2.5 to +9.5) values and206Pb/204Pb and 207Pb/204Pb isotope ratios in actinoliteschists as indication of the disturbance of the Sm–Nd andU–Pb isotope systems, rather than recording Mesoarch-ean mantle source heterogeneity.

Lead isotopes in pillow lavas (206Pb/204Pb=13.075–17.866; 207Pb/204Pb=14.025–15.425) display largerscatter than those in gabbros and diorites (206Pb/204

Pb=13.100–15.877; 207Pb/204Pb=14.027–14.687)(Table 8). This variation correlates well with the higherdegree of metasomatic alteration in the former group.

Although samples for this study were collected from theleast metasomatized pillow cores, some of these corescontain concentric cooling cracks filled with quartz andplagioclase (Fig. 3b). On the basis of Mesoarchean(3058±240 Ma) Pb–Pb errorchron age, we suggest thatmobilization of U and Pb took place during seafloorhydrothermal alteration, as recorded by mineralogicaland chemical zonation in pillow basalts (Figs. 3, 11).

The initial εNd values in gabbros and diorites (+0.30to +2.86) overlap with, but extend to lower values than,the pillow lavas (+1.10 to +3.10). We suggest that theseinitial εNd values likely reflect the near-primary magmaticcompositions for the following reasons: (1) both groupsplot co-linearly in a 143Nd/144Nd versus147Sm/144Nddiagram, yielding a Mesoarchean age (3060±220 Ma,MSWD=80); (2) there are no co-variations between theinitial εNd values and La/Smcn ratios within each group;and (3) there are no correlations between εNd and mobileelements (e.g. LILE) (Table 6).

6.2. Geodynamic setting

Stage I metasomatic assemblage in the Ivisaartoqgreenstone belt is analogous to that of Phanerozoicsupra-subduction ophiolites and intra-oceanic islandarcs (see Polat et al., 2007, Python et al., 2007, andreferences therein). The LREE-enriched trace elementpatterns of the Ivisaartoq volcanic and intrusive rocks,and those of anorthositic inclusions in gabbros (Figs. 8–10) are consistent with a subduction zone geochemicalsignature (cf. Saunders et al., 1991; Hawkesworth et al.,1993). However, a backarc origin for the belt cannot beruled out on the basis of geochemical data alone.

High MgO (14–24 wt.%), Ni (600–850 ppm) and Cr(1600–1700 ppm) concentrations, and Mg-numbers(70–83) in clinopyroxene cumulates and high-Mgpillow lavas are consistent with an island arc picriticcomposition. Tertiary island arc picrites have been

Table 8U–Th–Pb isotope compositions of the Ivisaartoq actinolite schists, cumulates, pillow lavas, gabbros, and diorites

Sample # Rock type 206Pb/204Pb ±2σ 207Pb/204Pb ±2σ 208Pb/204Pb ±2σ r1 r2 Th⁎ U⁎ Pb⁎ U/Pb Th/Pb

485426 Actinolite schist 15.645 0.030 14.591 0.029 34.931 0.075 0.98 0.91 0.155 0.264 1.023 0.258 0.152485430 Actinolite schist 15.590 0.011 14.590 0.012 34.594 0.032 0.96 0.94 0.121 0.444 0.924 0.481 0.130485431 Actinolite schist 12.687 0.010 13.870 0.012 32.355 0.032 0.96 0.93 0.369 0.134 1.862 0.072 0.198485433 Actinolite schist 14.287 0.013 14.373 0.014 33.910 0.036 0.97 0.94 0.060 0.023 5.162 0.004 0.012485434 Actinolite schist 13.575 0.008 14.040 0.010 32.788 0.027 0.96 0.92 0.064 0.079 5.595 0.014 0.011485435 Actinolite schist 13.760 0.021 14.157 0.023 32.898 0.056 0.98 0.95 0.193 0.035 8.567 0.004 0.023485436 Actinolite schist 15.617 0.011 14.399 0.011 33.783 0.030 0.96 0.93 0.085 0.084 2.325 0.036 0.037485437 Actinolite schist 12.854 0.009 13.892 0.011 32.393 0.029 0.96 0.93 0.085 0.084 5.970 0.014 0.014485473 Cpx cumulate 21.150 0.057 15.502 0.043 39.975 0.111 0.98 0.98 0.339 0.293 1.554 0.188 0.218485474 Cpx cumulate 19.656 0.025 15.325 0.021 38.868 0.059 0.97 0.88 0.259 0.668 1.652 0.405 0.157485475 Cpx cumulate 19.345 0.021 15.225 0.018 40.474 0.050 0.97 0.95 0.356 0.183 1.833 0.100 0.194485481 Pillow lava 17.816 0.015 15.069 0.014 37.301 0.039 0.95 0.91 0.683 0.148 1.897 0.078 0.360485482 Pillow lava 18.352 0.017 14.977 0.015 37.038 0.042 0.97 0.93 0.825 0.286 1.280 0.223 0.645485486 Pillow lava 17.693 0.028 14.852 0.025 38.322 0.065 0.98 0.97 0.722 0.260 2.776 0.094 0.260485414 Pillow lava 13.075 0.007 14.025 0.010 33.357 0.027 0.96 0.93 0.370 0.064 1.809 0.035 0.205485418 Pillow lava 13.741 0.015 14.155 0.018 33.696 0.048 0.91 0.84 0.402 0.101 2.229 0.045 0.180485420 Pillow lava 17.866 0.016 15.425 0.015 37.523 0.040 0.97 0.93 0.708 0.192 1.712 0.112 0.414485422 Pillow lava 17.603 0.017 14.961 0.015 36.254 0.040 0.97 0.95 0.552 0.167 1.257 0.133 0.439485468 Pillow lava 16.985 0.013 15.219 0.013 36.639 0.034 0.97 0.94 0.787 0.190 5.319 0.036 0.148485469 Pillow lava 15.226 0.020 14.666 0.020 35.167 0.050 0.98 0.97 0.478 0.183 3.196 0.057 0.149485428 Gabbro 14.917 0.012 14.551 0.013 35.237 0.035 0.97 0.94 0.217 0.035 1.046 0.033 0.208485432 Gabbro 13.521 0.014 14.111 0.016 33.190 0.039 0.97 0.96 0.515 0.188 4.441 0.042 0.116485438 Gabbro 13.320 0.006 14.133 0.009 33.123 0.025 0.97 0.93 0.532 0.100 6.033 0.017 0.088485467 Gabbro 13.958 0.007 14.264 0.009 33.453 0.025 0.97 0.94 0.726 0.160 3.463 0.046 0.210485472 Gabbro 13.100 0.008 14.027 0.011 32.923 0.029 0.95 0.91 0.465 0.149 4.399 0.034 0.106485476 Gabbro 13.238 0.008 14.096 0.010 33.158 0.028 0.96 0.92 0.628 0.176 6.433 0.027 0.098485477 Gabbro 13.516 0.011 14.152 0.013 33.888 0.034 0.96 0.94 0.503 0.257 7.461 0.034 0.067499730-A Gabbro 15.877 0.021 14.687 0.021 35.015 0.052 0.97 0.95 0.34 0.09 1.411 0.065 0.244499739 Diorite 14.355 0.008 14.349 0.009 34.183 0.026 0.95 0.93 0.793 0.185 4.934 0.037 0.161485483 Diorite 14.250 0.013 14.338 0.015 34.478 0.038 0.96 0.95 1.161 0.355 5.357 0.066 0.217485429 Diorite 15.138 0.015 14.587 0.015 35.006 0.039 0.98 0.95 0.686 0.150 2.543 0.059 0.270485484 Diorite 14.230 0.012 14.330 0.014 34.324 0.037 0.96 0.91 0.942 0.298 3.831 0.078 0.246

r1=206Pb/204Pb vs. 207Pb/204Pb error correlation (Ludwig, 1988).

r2=206Pb/204Pb vs. 208Pb/204Pb error correlation (Ludwig, 1988).

Errors are two standard deviations absolute (Ludwig, 1988).Fractionation of Pb was controlled by repeated analysis of the SRM NBS 981 standard (values of Todt et al., 1993) and amounted to 0.103+/−0.007%/a.m.u. (2 s, n=5).*U, Th, Pb concentrations represent measurements by high-resolution ICP-MS at Windsor.


documented in the Solomon and NewHebrides (Vanuatuarc) oceanic island arcs (see Eggins, 1993; Schuth et al.,2004). In the Solomon Islands, picrites occur only inNew Georgia Island above the subducting Woodlarkspreading centre. In the New Hebrides subductionsystem, the overriding plate is currently undergoingextension east of the Vanuatu arc, forming a supra-subduction oceanic crust within the North Fiji Basin(Hawkins, 2003). As a corollary, given that thegeochemical characteristics and hydrothermal alterationfeatures of the Ivisaartoq rocks are comparable to thoseof Phanerozoic ophiolites and intra-oceanic island arcs,we suggest that the Ivisaartoq belt originated asMesoarchean supra-subduction oceanic crust (Fig. 12).

6.3. Petrogenesis and mantle source characteristics

The positive initial εNd values (e.g. +4.2 to +5.0 inclinopyroxene cumulates; +4.8 to +6.0 in anorthositicinclusions; +1.1 to +3.1 in pillow basalts) require long-term depleted upper mantle sources. Given near-flatHREE patterns in these lithologies, melting occurred atb80 km (cf. Hirschmann and Stolper, 1996). Picriticcumulates and anorthositic inclusions plot above theestimated evolution curve of the depletedmantle (Fig. 13).The depleted Nd isotopic signatures and low LREE andHFSE (Nb, Ta, Zr, Ti, Y) abundances indicate that themantle source region had experienced significant meltextraction prior to 3075 Ma.

Fig. 12. A simplified geodynamic and petrologic model for the Ivisaartoq greenstone belt and Mesoarchean forearc oceanic crust.


The majority of pillow lavas, gabbros, and dioritesplot below the predicted depleted mantle evolution curve(Fig. 13; cf. Henry et al., 2000; Bennett, 2003). Largevariations in the initial εNd (+0.30 to +3.10) values mayreflect either mantle source heterogeneity or crustalcontamination. Contamination of the Ivisaartoq rocks bycontinental crust during magma ascent, rather thancontamination of their source regions by subductedcrustal material, can be ruled out on the basis of thefollowing observations: (1) the association of pillow

basalts, gabbros, and sulphide-rich siliceous volcani-clastic sedimentary rocks is expected to have formed inan oceanic rather than a continental setting; (2) there isno field evidence indicating that the Ivisaartoq green-stone belt was deposited on an older continentalbasement; (3) the lack of co-variations between εNdand abundances of contamination-sensitive elements ortheir ratios (e.g. SiO2, Ni, Cr, Th, Zr, La/Smcn) withineach volcanic group (Tables 4, 6, 7); and (4) there are nocorrelation between 147Sm/144Nd initial εNd values,

Fig. 13. Age versus initial εNd variation diagram for the Ivisaartoq cumulates, pillow lavas, gabbros, and diorites, Wawa adakites and Mg-andesites(Polat and Kerrich, 2002), Kostomuksha komatiites (Puchtel et al., 1998), and Winnipeg River granitoids (Henry et al., 2000). Modified after Henryet al. (2000). (DM: Depleted mantle; CHUR: Chondrite uniform reservoir).


which, according to Vervoort and Blichert-Toft (1999), isa robust criterion for identifying crustal contamination(Table 6). Accordingly, the lower initial εNd values(b+2.0) likely indicate a Nd-enriched component in thesource region, rather than crustal contamination. Wetherefore suggest that an enriched component was addedto the mantle wedge in variable proportions by recyclingof older continental material, with super-chondritic Nd/Sm ratios (cf. Shirey and Hanson, 1986; Polat andKerrich, 2002).

Light-REE-enriched patterns of the Ivisaartoq rocks,however, imply that the depleted sub-arc mantle sourcemust have been metasomatized shortly before or duringpartial melting that took place at about 3075 Ma (cf. Polatand Münker, 2004). Hydrous fluids and/or melts derivedfrom either subducted altered oceanic crust or sediments,with sub-chondritic Nb/La, Nb/Th, Sm/Nd, and Ti/Gdratios, were probably the main cause of themetasomatism,generating LREE-enriched, HFSE-depleted trace elementpatterns in the Ivisaartoq rocks (Figs. 8, 10).

6.4. Origin of the anorthositic inclusions in gabbros

Ocellar texture characterized by eye-shaped anorthositeinclusions and gabbroic matrix (Figs. 4f, 5a) was pre-viously interpreted as solidified immiscible liquids (Polatet al., 2007). However, new petrographic evidence, such asthe presence of relic reaction rims and smaller pieces ofpeeled anorthosites at the inclusion-matrix contacts, sug-gest that the inclusions had already been solidified beforethey were enclosed in gabbroic magma. The smalleranorthositic inclusions, aligned parallel to the contacts(Fig. 5a), might have resulted from the thermal erosion ofthe larger ones during their transportation to shallowerdepths. Therefore, we suggest that the anorthositic inclu-sions were carried, as crustal xenoliths, from lower oceanic

crust to the shallower depths by upwelling magmas. Boththe inclusions andmatrix were deformed and recrystallizedunder amphibolite facies metamorphic conditions beforethe intrusion of 2961±12 Ma granitoids.

Low abundances of MgO, Ni, Cr, Co, and Sc in theanorthositic inclusions are consistent with olivine, clin-opyroxene and/or orthopyroxene fractionation (Fig. 10;Table 5). Depletion of Nb, relative to Th and La, and near-flat HREE patterns are consistent with a subduction zonegeodynamic setting and a shallow mantle source (Figs. 10,12; Table 5). The gabbroic matrix shares the negative Nbanomalies (Fig. 10; Table 5). The initial εNd values of theanorthositic inclusions are much larger than those of thegabbroic matrix (+4.8 to +6.0 versus +2.3 to +2.8),indicating two different mantle sources. On the other hand,the anorthositic inclusions are isotopically comparable topicrites (clinopyroxene cumulates) (Tables 6, 7), indicatinga petrogenetic link between the two rock types. It is likelythat picrites and anorthosites were derived from the sameparental magma through clinopyroxene and plagioclasefractionation, respectively. Given their low viscosity,picritic magmas could easily have reached the surface toform the ultramafic sills and/or flows. In contrast,anorthositic magmas might have been too viscous toreach to the surface; instead, they might have crystallizedwithin the lower oceanic crust to form layered anorthosites.

7. Implications for the generation of Mesoarcheanoceanic crust

Field relationships and geochemical data indicate thatall volcanic and intrusive rock types in the Mesoarchean(ca. 3075 Ma) Ivisaartoq greenstone belt are part of thesame lithotectonic assemblage, sharing a common historyof magmatism, deformation, and metamorphism. Theassociation of pillow basalt, gabbros and sulphide-rich


siliceous volcaniclastic sedimentary rocks in the beltsuggests an intra-oceanic depositional environment(Figs. 3–5).

The large initial εNd (e.g. +2 to +6) isotopic values inthe Mesoarchean Ivisaartoq rocks indicate a long-termLREE-depleted (Sm/NdcnN1) mantle source(s), similar tothe source of modern N-MORB (see Hofmann, 2003).However, the majority of the least altered samples haveLREE-enriched (La/SmcnN1; Sm/Ndcnb1) but Nb-deplet-ed, relative to Th and La, trace element patterns (Figs. 9,10), consistent with a subduction zone geodynamic setting(Fig. 12). Accordingly, we propose a two-stage evolution-ary geodynamic model to explain the geological character-istics of the Ivisaartoq greenstone belt. In the first stage, themantle source of the Ivisaartoq rocks had originated as asub-oceanic depleted upper mantle, like the source ofpresent-day N-MORB. The second stage marks thedevelopment of an intra-oceanic subduction system.Following the initiation of an intra-oceanic subductionzone along either amid-ocean ridge or a transform fault, themantle source of the Ivisaartoq rocks was converted to asub-arc mantle wedge (cf. Casey and Dewey, 1984; Dilekand Flower, 2003). Hydrous fluids and/or melts originatingfrom the subducted slab metasomatized the sub-arc mantlewedge, resulting in LREE-enriched and HFSE-depleted,relative to Th and LREE, patterns (Figs. 8, 10).

The forearc region of the overriding plate may haveundergone a significant extension in response to slabrollback, resulting in a large degree of partial melting of thehydrated upper mantle wedge at shallow depths (cf. Dilekand Flower, 2003; Fig. 12). Such a high degree of partialmelting is expected to have resulted in the formation of alarge magma chamber (Fig. 12). Extensive partial meltingbeneath the Ivisaartoq forearc may have generated a thick(N20 km) oceanic crust (cf. Sleep and Windley, 1982).Such an intact oceanic crust might have been composed oftwomajor crustal sections: (1) a lower layer of anorthositesand leucogabbros; and (2) an upper layer of basaltic topicritic flows, gabbroic to dioritic dykes, and dunitic towehrlitic sills (Figs. 2, 12).

The major and trace element characteristics of theanorthositic inclusions in the Ivisaartoq are comparableto those of Meso- to Neoarchean anorthosite complexesin SW Greenland (Windley et al., 1973; Weaver et al.,1981; Ashwal and Myers, 1994; Dymek and Owens,2001). Like the Buksefjorden, Nordland, and Fiskenaes-set anorthosites, the Ivisaartoq counterparts have LREEenriched chondrite-normalized patterns with large pos-itive Eu anomalies (see Weaver et al., 1981; Dymek andOwens, 2001), suggesting a similar petrogenetic process.However, the Ivisaartoq anorthositic inclusions have flatto less fractionated HREE patterns compared to the

Buksefjorden, Nordland, and Fiskenaesset anorthosites,indicating a shallower, garnet-free mantle source region(see Dymek and Owens, 2001).

In the Ivisaartoq belt, anorthosites and leucogabbrosoccur as volumetrically minor intrusions within the loweramphibolites (Figs. 1, 2; Chadwick, 1990); however, theymight originally have been thicker. There are two mainreasons why this might be the case. First, given the recordof several generations of deformation in the region(Chadwick, 1990; Friend and Nutman, 2005), finding anintact, thicker leucogabbro–anorthosite association isunlikely. Second, if Archean oceanic crust was thickerdue to potentially higher mantle temperatures (Sleep andWindley, 1982; McKenzie and Bickle, 1988), then it ispossible that only the upper basaltic crustal section waspeeled off and accreted while the lower anorthosite–leucogabbro section was subducted. Notwithstandingthese problems, partial sections of 2800–3000 Maanorthosite–leucogabbro associations have been identi-fied throughout the Archean terranes of southern WestGreenland (Windley, 1970; Windley et al., 1973; Escherand Myers, 1975; Windley et al., 1981; Myers, 1985;Ashwal and Myers, 1994; Owens and Dymek, 1997).

The petrogenetic origin of Archean anorthositesremains unresolved (Weaver et al., 1981, 1982; Pinneyet al., 1988; Ashwal and Myers, 1994; Owens andDymek, 1997; Dymek and Owens, 2001). In the best-studied Fiskenaesset anorthosite complex, the anortho-sites and gabbros appear to have intruded into theoverlying greenstone sequences (Escher and Myers,1975; Ashwal and Myers, 1994). Geochemical studiessuggest that the Fiskenaesset anorthosite complex isgenetically related, by fractional crystallization, to maficto ultramafic volcanic rocks into which they wereemplaced (Weaver et al., 1981, 1982; Peck and Valley,1996) and were derived from a long-term depletedmantle source (Ashwal et al., 1989). The anorthositicinclusions in gabbros, and anorthosites layers intrudingthe lower amphibolites, in the Ivisaartoq belt are likelyrelated to the ultramafic lithologies in the belt byfractional crystallization.

Greenstone–anorthosite associations in SWGreenland(e.g. Fiskenaesset Complex) are interpreted as remnantArchean oceanic crust (ophiolite?) in several studies(Windley et al., 1981; Weaver et al., 1982; Myers, 1985;Ashwal and Myers, 1994). On the basis of geologicalsimilarities between the Ivisaartoq and Fiskenaessetgreenstone belts, and geological and geochemical datapresented in this study, we interpret the Ivisaartoqgreenstone belt as a relic of an Archean forearc oceaniccrust. We do not propose that all Archean anorthositesformed in a forearc tectonic setting. Like basaltic


counterparts, Archean anorthosites, depending on theirgeological and geochemical characteristics, might haveformed in diverse geodynamic settings, including in mid-ocean ridges, forearcs, backarcs, plume-derived oceanicplateaus, and intra-continental rifts.

Acknowledgements

We thank A. Trenhaile and R. Kerrich for reviewingthe initial draft of the manuscript. J.C. Barrette, and J.Gagnon are acknowledged for their help during geo-chemical analyses. Reviewers J. Dostal, H. Smithies, P.Spadea, and M.F. Zhou are acknowledged for theirconstructive comments, which have resulted in significantimprovements to the paper. We are grateful to B.F.Windley for helpful discussion on the geology of Archeangreenstone belts and anorthosite complexes. This is acontribution ofNSERCgrants 250926 toAP and 83117 toB. Fryer. R. Frei is supported by FNU (Forskningsrådetfor Natur of Univers) grant no. 21-01-0492 56493. Fieldwork was supported by the Bureau of Minerals andPetroleum in Nuuk and the Geological Survey ofDenmark and Greenland (GEUS). A. Polat thanks Mr.Tekin Demir for helping his family during fieldwork inGreenland.

References

Abbott, D., Drury, R., Smith, W.H.F., 1994. Flat to steep transition insubduction style. Geology 22, 937–940.

Appel, P.W.U., 1994. Stratabound scheelite in altered Archeankomatiites, West Greenland. Miner. Depos. 29, 341–352.

Appel, P.W.U., 1997. High bromine and low Cl/Br ratios inhydrothermally altered Archean komatiitic rocks, West Greenland.Precambrian Res. 82, 177–189.

Arndt, N.T, Bruzak, G., Reischmann, T., 2001. The oldest continentaland oceanic plateaus-geochemistry of basalts and komatiites of thePilbara Craton,Australia. In: Buchan, K.L, Ernst, R.E. (Eds.),Mantleplumes: their identification through time. Geol. Soc. Am. Spec.Paper, vol. 352, pp. 1–30.

Ashwal, L.D., Myers, J., 1994. Archean anorthosites. In: Condie,K.C. (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam,pp. 315–355.

Ashwal, L.A., Jacobsen, S.B., Myers, J.S., Kalsbeek, F., Goldstein, S.J.,1989. Sm–Nd age of the Fiskenaesset Anorthosite Complex, WestGreenland. Earth Planet. Sci. Lett. 91, 261–270.

Bennett, V.C., 2003. Compositional evolution of the mantle. In: Carlson,R.V. (Ed.), Treatise on Geochemistry, vol. 3. Elsevier, Amsterdam,pp. 493–519.

Bickle, M.J., Nisbet, E.G., Martin, A., 1994. Archean greenstone beltsare not oceanic crust. J. Geol. 102, 121–138.

Bleeker, W., 2002. Archean tectonics: a review, with illustrationsfrom the Slave craton. In: Fowler, M.R., Ebinger, C.J., Hawkes-worth, C.J. (Eds.), The Early Earth: Physical, Chemical andBiological Development. Geol. Soc. London Spec. Publ., vol. 199,pp. 151–181.

Brewer, M.A., Chadwick, B., Coe, K., Park, J.F.W., 1984. Further fieldinvestigations in the Ivisaartoq region of southern West Greenland.Rapp. - Grùnl. Geol. Unders. 120, 55–67.

Casey, J.F., Dewey, J.F., 1984. Initiation of subduction zones alongtransform and accreting plate boundaries, triple-junction evolution,and forearc spreading centres— implications for ophiolitic geologyand obduction. In: Gass, I.G., Lippard, S.J., Shelton, A.W. (Eds.),Ophiolites and Oceanic Lithosphere. Geol. Soc. London, Spec.Publ., vol. 13, pp. 269–290.

Chadwick, B., 1985. Contrasting styles of tectonism and magmatismin the late Archean crustal evolution of the northeastern part of theIvisaartoq region, inner Godthåbsfjord, southern West Greenland.Precambrian Res. 27, 215–238.

Chadwick, B., 1986. Malene stratigraphy and late Archean structure:new data from Ivisaartoq, inner Godthåbsfjord, southern WestGreenland. Rapp. - Grùnl. Geol. Unders. 130, 74–85.

Chadwick, B., 1990. The stratigraphy of a sheet of supracrustal rockswithin high-grade orthogneisses and its bearing on late Archeanstructure in southern West Greenland. J. Geol. Soc. (Lond.) 147,639–652.

Chadwick, B., Coe, K., 1988. Geological Map of Greenland,1:100000, Ivisaartoq Sheet 64 V. Nord. Copenhagen, GeologicalSurvey of Greenland.

Cloos, M., 1993. Lithospheric buoyancy and collisional orogenesis:subduction of oceanic plateaus, continental margin, island arcs,spreading ridges, and seamounts. GSA Bull. 105, 715–737.

Condie, K.C., 1981. Archean Greenstone Belts. Elsevier, Amsterdam.435 pp.

Condie, K.C., 2005. Earth as an Evolving Planetary System. Elsevier,Amsterdam. 447 pp.

Corcoran, P.L., Dostal, J., 2000. Development of an ancient-back arcbasin overlying continental crust: the Archean Peltier Formation,Northwest Territories, Canada. J. Geol. 109, 329–348.

Dewey, J., 2003. Ophiolites and lost oceans: rifts, ridges, arcs, and/orscraping. In: Dilek, Y., Newcomb, S. (Eds.), Ophiolite Concept andthe Evolution of Geological Thought. Geol. Soc. Am. Spec. Paper,vol. 373, pp. 153–158.

Dilek, Y., 2003. Ophiolite concept and its evolution. In: Dilek, Y.,Newcomb, S. (Eds.), Ophiolite Concept and the Evolution ofGeological Thought. Geol. Soc. Am. Spec. Paper, vol. 373, pp. 1–15.

Dilek, Y., Flower, M.F.J., 2003. Arc-trench rollback and forearcaccretion: 2. A model template for ophiolites in Albania, Cyprus,and Oman. In: Dilek, Y., Robinson, P.T. (Eds.), Ophiolites in EarthHistory. Geol. Soc. London Spec. Publ., vol. 218, pp. 43–68.

Dostal, J., Mueller, W.U., 1997. Komatiite flooding of a rifted Archeanrhyolite arc complex: geochemical signature and tectonic signif-icance of Stoughton-Roquemaure Group, Abitibi greenstone belt,Canada. J. Geol. 105, 545–563.

Dostal, J., Mueller, W.U., Murphy, J.B., 2004. Archean molasses basinevolution and magmatism, Wabigoon subprovince, Canada. J. Geol.112, 435–454.

Dymek, R.F., Owens, B.R., 2001. Chemical assembly of Archeananorthosites from amphibolite- and granulite-facies terranes, SWGreenland. Contrib. Mineral. Petrol. 141, 513–528.

Eggins, S.M., 1993. Origin and differentiation of picritic arcmagmas, Ambae [Aoba], Vanuatu. Contrib. Mineral. Petrol.114, 79–100.

Eriksson, K.A., Krapez, B., Fralick, P.W., 1994. Sedimentology ofArchean greenstone belts: signatures and tectonic evolution. EarthSci. Rev. 37, 1–88.

Escher, J.C., Myers, J.S., 1975. New evidence concerning the originalrelationships of early Precambrian volcanics and anorthosites in


the Fiskenaesset region, southern West Greenland. Rapp. - Grùnl.Geol. Unders. 75, 72–76.

Frei, R., Villa, I.M., Nägler, Th.F., Kramers, J.D., Pryzbylowicz, W.J.,Prozesky, V.M., Hofmann, B.A., Kamber, B.S., 1997. Singlemineral dating by Pb–Pb step-leaching method; Assessing themechanisms. Geochim. Cosmochim. Acta 61, 393–414.

Friend, C.R.L., Nutman, A.P., 2005. New pieces to the Archean jigsawpuzzle in the Nuuk region, southern West Greenland: steps intransforming a simple insight into a complex regional tectonother-mal model. J. Geol. Soc. (Lond.) 162, 147–162.

Friend, C.R.L., Hall, R.P., Hughes, D.J., 1981. The geochemistry of theMalene (mid-Archean) ultramafic–mafic amphibolite suite, south-ern West Greenland. In: Glover, J.E., Groves, D.E. (Eds.), ArcheanGeology. Spec. Pub. Geol. Soc. Australia, vol. 7, pp. 301–312.

Goodwin, M.A., 1991. Precambrian Geology, The Dynamic Evolutionof the Continental Crust. Academic press, London. 666 pp.

Hall, R.P., 1981. The tholeiitic and komatiitic affinity of Malenemetavolcanic amphibolites from Ivisaartoq, southern WestGreenland. Rapp. - Grùnl. Geol. Unders. 97 20 pp.

Hall, R.P., Friend, C.R.L., 1979. Structural evolution of the Archeanrocks in Ivisaartoq and neighbouring inner Godthåbsfjord region,southern West Greenland. Geology 7, 311–315.

Hall, R.P., Hughes, D.J., Friend, C.R.L., 1987. Mid-Archean basicmagmatism of southern West Greenland. In: Park, R.G., Tarney, J.(Eds.), Evolution of the Lewisian and Comparable PrecambrianHigh Grade Terrains. Geol. Soc. London Spec. Publ., vol. 27,pp. 261–275.

Hawkesworth, C.J., Gallanger, K., Hergt, J.M., McDetmott, F., 1993.Mantle and slab contributions in arc magmas. Annu. Rev. EarthPlanet. Sci. 21, 175–204.

Hawkins, J.W., 2000. Geology of supra-subduction zones—implica-tions for the origin of ophiolites. In: Dilek, Y., Newcomb, S. (Eds.),Ophiolite Concept and the Evolution of Geological Thought. Geol.Soc. Am. Spec. Paper, vol. 373, pp. 227–268.

Henry, P., Stevenson, R.S., Larbi, Y., Gariépy, C., 2000. Nd isotopicevidence for early to late Archean (3.4–2.7 Ga) crustal growth inthe Western Superior Province (Ontario, Canada). Tectonophys322, 135–151.

Hirschmann, M.M., Stolper, E.M., 1996. A possible role for garnetpyroxenite in the origin of “garnet signature” in MORB. Contrib.Mineral. Petrol. 124, 185–208.

Hofmann, A.W., 1988. Chemical differentiation of the Earth: therelationships between mantle, continental crust, and oceanic crust.Earth Planet. Sci. Lett. 90, 297–314.

Hofmann, A.W., 2003. Samplingmantle heterogeneity through oceanicbasalts: isotopes and trace elements. In: Carson, R.W. (Ed.),Treatise on Geochemistry, Th mantle and Core, vol. 2, pp. 61–101.

Kerrich, R., Polat, A., 2006. Archean greenstone–tonalite duality:thermochemical mantle convection models or plate tectonics in theearly Earth global dynamics? Tectonophysics 415, 141–165.

Kusky, T.M., 2004. Introduction. In: Kusky, T.M. (Ed.), PrecambrianOphiolites and Related Rocks. Elsevier, Amsterdam, pp. 1–34.

Kusky, T.M., Polat, A., 1999. Growth of granite–greenstone terranes atconvergent margins, and stabilization of Archean cratons. Tectono-physics 305, 43–73.

Ludwig, K.R., 1988. A computer program to convert raw U–Th–Pbisotope ratios to blank corrected isotope ratios and concentrationswith associated error-correlations: United States GeologicalSurvey. Open File Report. OF-82-820.

Ludwig, K., 2003. User's Manual for Isoplot/Ex Version 3.0, AGeochronological Toolkit for Microsoft Excel. Special Publication1a, Berkeley Geochronology Center.

McKenzie, D., Bickle, M.J., 1988. The volume and composition ofmelt generated by extension of the lithosphere. J. Petrol. 29,625–679.

Myers, J.S., 1985. Stratigraphy and structure of the FiskenaessetComplex,West Greenland. GronlandGeol. Unders. Bull. 150. 72 pp.

Nisbet, E.G., Fowler, C.M.R., 1983. Model for Archean platetectonics. Geology 11, 376–379.

Owens, B.E., Dymek, R.F., 1997. Comparative petrology of Archeananorthosites in amphibolite and granulite facies terranes, SWGreenland. Contrib. Mineral. Petrol. 128, 371–384.

Peck, W.H., Valley, J.W., 1996. The Fiskenaesset AnorthositeComplex: stable isotope evidence for shallow emplacement intoArchean oceanic crust. Geology 24, 523–526.

Pinney, W.C., Morrison, D.A., Maczuga, D.E., 1988. Anorthosites andrelated megacrystic units in the evolution of Archean crust.J. Petrol. 29, 1283–1323.

Polat, A., Hofmann, A.W., 2003. Alteration and geochemical patternsin the 3.7–3.8 Ga Isua greenstone belt, West Greenland.Precambrian Res. 126, 197–218.

Polat, A., Kerrich, R., 2002. Nd-isotope systematics of ∼2.7 Gaadakites, magnesian andesites, and arc basalts, Superior Province,Canada: evidence for shallow crustal recycling at Archeansubduction zones. Earth Planet. Sci. Lett. 202, 345–360.

Polat, A., Kerrich, R., 2006. Reading the geochemical fingerprints ofArchean hot subduction volcanic rocks: evidence for accretion andcrustal recycling in amobile tectonic regime. In: Benn,K.,Mareschal,J.C., Condie, K.C. (Eds.), Archean Geodynamics and Environments.AGU Geophys. Monograph Series, vol. 164, pp. 189–213.

Polat, A., Münker, C., 2004. Hf-Nd isotope evidence for contempo-raneous subduction processes in the source of late Archean arclavas from the Superior Province, Canada. Chem. Geol. 213,403–429.

Polat, A., Kerrich, R.,Wyman, D.A., 1998. The late Archean Schreiber-Hemlo and White River-Dayohessarah greenstone belts, SuperiorProvince: collages of oceanic plateaus, oceanic arcs, andsubduction–accretion complexes. Tectonophysics 294, 295–326.

Polat, A., Appel, P.W.U., Frei, R., Pan, Y., Dilek, Y., Ordonez-Caldeeron, J.C., Fryer, B., Raith, J.G., 2007. Field andGeochemical Characteristics of the Mesoarchean (∼3075 Ma)Ivisaartoq Greenstone Belt, Southern West Greenland: Evidencefor Seafloor Hydrothermal Alteration in a Supra-SubductionOceanic Crust. Gondwana Research 11, 69–91.

Puchtel, I.S., Hofmann, A.W., Mezger, K., Jochum, K.P., Shchipansky,A.A., Samsonov, A.V., 1998. Oceanic plateau model forcontinental crustal growth in the Archean: a case study from theKostomuksha greenstone belt, NW Baltic Shield. Earth. Planet.Sci. Lett. 155, 57–74.

Python, M., Ceuleneer, G., Ishida, Y., Barrat, J.A., Arai, S., 2007. Omandiopsides: a new lithology diagnostic of very high temperaturehydrothermal circulation in mantle peridotite below oceanicspreading centres. Earth Planet. Sci. Lett. 255, 289–305.

Sandeman, H.A., Hanmer, S., Davis, W.J., Ryan, J.J., Peterson, T.D.,2004. Neoarchean volcanic rocks, Central Hearne supracrustalbelt, Western Churchill Province, Canada: geochemical andisotopic evidence supporting intra-oceanic, suprasubduction zoneextension. Precambrian Res. 134, 113–141.

Saunders, A.D., Norry, M.J., Tarney, J., 1991. Fluid influence on thetrace element compositions of subduction zone magmas. Phil.Trans. Royal Soc. London A 335, 377–392.

Schaller, M., Steiner, O., Suder, I., Frei, R., Kramers, J.D., 1997. Pbstepwise leaching (PbSL) dating of garnet-addressing the inclusionproblem. Schweiz. Mineral. Petrogr. Mitt. 77, 113–121.


Schuth, S., Rohrbach, A., Münker, C., Ballhaus, C., Garbe-Schönberg,D., Qopoto, C., 2004. Geochemical constraints on the petrogenesisof arc picrites and basalts, New Georgia Group, Solomon Islands.Contrib. Mineral. Petrol. 148, 288–304.

Şengör, A.M.C., Natal'in, B.A., 2004. Phanerozoic analogues ofArchean oceanic basement fragments: Altaid ophiolites andophirags. In: Kusky, T.M. (Ed.), Precambrian Ophiolites andRelated Rocks. Elsevier, Amsterdam, pp. 675–726.

Shirey, S.B., Hanson, G.N., 1986. Mantle heterogeneity and crustalrecycling in Archean granite–greenstone belts: Evidence from Nd-isotopes and trace elements in the Rainy Lake area, Superior Province,Ontario, Canada. Geochim. Cosmochim. Acta 50, 2631–2651.

Sleep, N.H., Windley, B.F., 1982. Archean plate tectonics: constrainsand inferences. J. Geol. 90, 363–379.

Smithies, R.H., Champion, D.C., van Kranendonk, M.J., Howard, H.M.,Hickman, A.H., 2005a. Modern-style subduction processes in theMesoarchean: geochemical evidence from the 3.12 Ga Whundointra-oceanic arc. Earth Planet. Sci. Lett. 231, 221–237.

Smithies, R.H., van Kranendonk, M.J., Champion, D.C., 2005b. Itstarted with a plume — early Archean basaltic proto-continentalcrust. Earth Planet. Sci. Lett. 238, 284–297.

Sun, S.S.,McDonough,W.F., 1989. Chemical and isotopic systematics ofoceanic basalts: implications for mantle composition and processes.In: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the OceanBasins. Geol. Soc. London Spec. Publ., vol. 42, pp. 313–345.

Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: ItsComposition and Evolution. Blackwell, Oxford. 312 pp.

Todt, W., Cliff, R.A., Hanser, A., Hofmann, A.W., 1993. Recalibrationof NBS lead standards using a 202Pb+205Pb double spike. TerraAbstr. 5 (Suppl. 1), 396.

van Hunen, J., van den Berg, A.P., Vlaar, N.J., 2002. On the role ofoceanic plateaus in the development of shallow flat subduction.Tectonophysics 352, 317–333.

Vervoort, J.D., Blichert-Toft, J., 1999. Evolution of the depletedmantle: Hf isotope evidence from juvenile rocks through time.Geochim. Cosmochim. Acta 63, 533–556.

Weaver, B.L., Tarney, J., Windley, B., 1981. Geochemistry andpetrogenesis of the Fiskenaesset anorthosites complex, southernWest Greenland: Nature and parent magma. Geochim. Cosmochim.Acta 45, 711–725.

Weaver, B.L., Tarney, J., Windley, B., Leake, B.E., 1982. Geochem-istry and petrogenesis of Archean metavolcanic amphibolites fromFiskenaesset, S.W. Greenland. Geochim. Cosmochim. Acta 46,2203–2215.

Windley, B.F., 1970. The stratigraphy of the Fiskenaesset complex.Rapp. - Grùnl. Geol. Unders. 35, 19–21.

Windley, B.F., Herd, R.K., Bowden, A.A., 1973. The FiskenaessetComplex, West Greenland, Part 1. Grønlands Geol. Unders. Bull.106 80 pp.

Windley, B.F., Bishop, F.C., Smith, J.V., 1981. Metamorphosedlayered igneous complexes in Archean granulite–gneiss belts.Annu. Rev. Earth Planet. Sci. 9, 175–198.

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