Isua supracrustal belt (Greenland)—A vestige of a 3.8 Ga ... · The UA contains all major...

Lithos 113 (2009) 115–132

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Isua supracrustal belt (Greenland)—A vestige of a 3.8 Ga suprasubduction zoneophiolite, and the implications for Archean geology

Harald Furnes a,⁎, Minik Rosing b, Yildirim Dilek c, Maarten de Wit d

a Department of Earth Science & Centre for Geobiology, University of Bergen, Norwayb Nordic Centre for Earth Evolution, and Geological Museum, Univ. of Copenhagen, Denmarkc Department of Geology, Miami University, Oxford, OH, USAd AEON & Department of Geological Sciences, University of Cape Town, South Africa

⁎ Corresponding author. Department of Earth ScienceE-mail address: [email protected] (H. Furnes

0024-4937/$ – see front matter © 2009 Elsevier B.V. Aldoi:10.1016/j.lithos.2009.03.043

a b s t r a c t
a r t i c l e i n f o
Article history:Received 15 May 2008Accepted 20 March 2009Available online 6 April 2009

Keywords:Isua supracrustal beltOphiolitesSuprasubduction zone magmatismArchean oceanic crust and tectonics

The mafic–ultramafic units of the ∼3.8 Ga Isua supracrustal belt (ISB) in Greenland occur in a two-armedarcuate zone (eastern and western arms) and are grouped into two major tectonostratigraphic units basedon their lithological and geochemical characteristics: (1) Undifferentiated amphibolites (UA), and(2) Garbenschiefer amphibolites (GA). The UA contains all major lithological units of a typical Penrose-type complete ophiolite sequence. The GA is composed dominantly of volcaniclastic and volcanic rocks,commonly found in immature island arcs. The available geochemical data from UA and GA show distinctdifferences between the two units. Compared with the geochemical evolution of some of the well knownPhanerozoic ophiolites, the pillow lavas and associated dikes of the UA show a compositional range that issimilar to typical MORB-type Ligurian ophiolites in the Western Alps–Apennines and those displayed by LIP-type Caribbean ophiolites. The GA is characterized by island arc tholeiite (IAT) to boninite-like rocks anddefines a magmatic evolution that is comparable to that of suprasubduction zone (SSZ) ophiolites in theMediterranean region. Our proposed geodynamic model for the ISB suggests that the UAwas built by primaryto differentiated, mantle-generated melts during seafloor spreading, little to moderately affected bysubduction processes, and that the IAT to boninitic-like rocks of the GA formed at a later stage by meltingfrom a strongly subduction-affected, depleted and hydrated mantle. Our interpretation of the ISB is that theUA and GA represent early and late stages, respectively, in the formation of a SSZ ophiolite. This implies thatPhanerozoic-type plate tectonic processes, such as seafloor spreading and subduction, were operating by3.8 Ga in the Palaeoarchean.

© 2009 Elsevier B.V. All rights reserved.

1. Introduction

Archean supracrustal lithological associations, commonly referredto as greenstone belts, generally contain submarine mafic–ultramaficlavas and intrusions that vary compositionally from low- to high-MgObasalts to basaltic komatiites and komatiites (e.g. Arndt and Nesbitt,1982; de Wit and Ashwal, 1995, 1997; de Wit, 2004; Sproule et al.,2002). There is a general agreement among geologists working onArchean geology that a plate tectonic-like Earth, with divergent andconvergent plate boundaries, is consistent with field and geochemicaldata from the late Archean greenstone belts (b3.0 Ga; e.g. Goodwinand Ridler, 1970; Tarney et al., 1976; Condie, 1981; Windley, 1993; deWit, 1998; Kusky and Polat, 1999; Kerrich and Polat, 2006). However,estimates on precisely when this mode of horizontal tectonics was

, University of Bergen, Norway.).

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established vary by more than 1 billion years, e.g. at 3.2 Ga (vanKranendonck, 2007); at 3.6 Ga (Nutman et al., 2007); by 3.8 Ga(Furnes et al., 2007a,b; Dilek and Polat, 2008); at 4.0 Ga (de Wit,1998); by 4.2 Ga (Cavosie et al., 2007). Whereas some would arguethat plate tectonic processes operated throughout the Precambrian,others suggest that Phanerozoic-like plate tectonic processes did notcommence until the Neoproterozoic (Hamilton, 1998, 2003; Stern,2005; Brown, 2006). Some of the alternative interpretations envisagethat Archean tectonics was instead controlled by vertically controlledplume activities and associated crustal delamination.

One of the most critical criteria for modern plate tectonics is thesubduction-driven horizontal motion of lithospheric plates, resulting inchanges in their spatial relationship over time (e.g. Cawood et al., 2006).Palaeomagnetic studies now exist that demonstrate that large-scalehorizontal plate motions similar to those known from the PhanerozoicEon may have occurred in the Archean. Both examples are from thePilbara craton: the first is one from a late Archean supracrustal sequence(∼2.7 Ga by Strik et al., 2003) and the second from an early Archeansupracrustal sequence (3.4 Ga, by Suganuma et al., 2006). These results,

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116 H. Furnes et al. / Lithos 113 (2009) 115–132

however, are under some doubt as to their reliability because of theabsence of robust paleomagnetic field tests. Geological observations, onthe other hand, support the concept of early Archean horizontal crustalmotions both in terms of extension (e.g. de Vries et al., 2006) andshortening (e.g. deWit and Ashwal, 1997;Moyen et al., 2006), but noneof these inferred paleo-motions have been properly quantified. On theother hand, many geochemical studies of greenstone belts havedemonstrated similarities between Archean and modern mafic mag-matic rocks, which formed in different plate tectonic settings (e.g.Tarney et al., 1976; Lafleche et al., 1992; Kerrich et al., 1998; Kusky andPolat, 1999). One particularly debated question pertains to thesimilarities between Archean komatiites and Phanerozoic boninites,the latter being exclusively related to subduction zone environments(e.g. Arndt, 2003; Arndt et al., 1998, 2008; Parman et al., 2001, 2003;Parman and Grove, 2004; Grove and Parman, 2004; Dilek and Polat,2008).

Identification of ophiolites in the rock record has become one ofthe geological keystones with which to unravel plate tectonicprocesses in the rock record, because the majority of ophiolitic rocksuites are now widely recognized as sections of ancient oceanic crust/lithosphere formed in subduction rollback cycles (Dilek and Flower,2003). Therefore, recognition of ophiolites in the Archean record is ofutmost significance to establish the mode and tempo of plate tectonicprocesses in the early evolution of the Earth. The presence of Archeanophiolites, or ophiolite-like sequences, has been proposed by severalauthors for a large number of Archean greenstone belts (e.g. de Witet al., 1987; Kerrich et al., 1998; Kusky and Polat, 1999; Kusky et al.,2001; Polat et al., 2002; Polat and Hofmann, 2003; deWit, 2004; Dilekand Polat, 2008). However, Archean greenstone belts are commonlybelieved to lack one or more of the crustal components of the Penrosepseudostratigraphy (Anonymous, 1972), and this has led some

Fig. 1. Geological map of Isua. The compilation is mainly based on the detailed, regionalundifferentiated amphibolites (UA) are after Furnes et al. (2007a).

authors to conclude that ophiolites are not represented in the earlieststages of the Earth's history (Bickle et al., 1994; Hamilton, 1998, 2003).Thus, whether ophiolites as we know from the Phanerozoic rockrecord exist in the Archean greenstone belts remains a fundamentalquestion in global tectonics.

The Paleoarchean Isua supracrustal (ISB) or greenstone belt (e.g.Nutman et al., 1997) in southwestern Greenland (Fig. 1) contains oneof the oldest intact submarine igneous sequences in the world, andhence it presents an excellent opportunity to address the “Archeanophiolite and plate tectonic conundrum”. In a series of studies, basedon mapping and understanding of field relations, geochronologicaland geochemical studies of the volcanic and volcanogenic rocks, someresearchers have proposed that plate tectonic-like processes musthave operated during the generation of the ISB and its surroundingregion in the early Archean (Maruyama et al., 1994; Rosing et al., 1996;Komiya et al., 1999; Hanmer and Greene, 2002; Hanmer et al., 2002;Polat et al., 2002; Polat and Hofmann, 2003; Polat and Frei, 2005).Furnes et al. (2007a,b) identified a well-preserved sheeted dikecomplex within the suite of undifferentiated amphibolites andultramafic rocks of the western arm of the ISB. This observationimplies that the ISB has a highly deformed ophiolite that might haveformed as a result of plate tectonic-like processes around 3.7–3.8 Ga,although this interpretation has been disputed (Hamilton, 2007;Nutman and Friend, 2007). In this paper, we present additionalgeochemical data in support of our earlier interpretation that the ISBcontains ophiolite assemblages. We compare the geology andgeochemistry of the ISB to those of Phanerozoic ophiolites, whichformed in different tectonic environments at different stages of theirevolution. We then discuss the significance of the occurrence ofsuprasubduction zone (SSZ) ophiolites in the Archean record for theearly stages of the Earth's evolution.

map of Nutman (1986). The numbers on the eastern part of the western arm of the

117H. Furnes et al. / Lithos 113 (2009) 115–132

2. Geology of Isua

2.1. Rock types and tectonostratigraphic units

The ISB (Nutman et al., 1984) and the Akilia Association (McgregorandMason,1977; Baadsgaard et al., 1984) of mainly supracrustal rockswere formed in a similar time span, 3800–3500 Ma, named the IsuanEra by Harland et al. (1982). They are invariably older than the felsicAmitsoq orthogneisses (Mcgregor, 1973) that form part of the ItsaqGneiss Complex (Nutman et al., 1996). The exact age of the ISB andAkilia rocks are unknown (Nutman et al., 1993; Moorbath et al., 1997;Whitehouse et al., 1999; Nutman et al., 2000, 2007). The ISB (Fig. 1) iscomposed of a broad suite of lithologies that can be divided into threecategories: (1) rocks with preserved primary features that allowidentification of their protoliths, (2) rocks with ambiguous affinities,and (3) rocks that are demonstrably metasomatic in origin. The firstcategory is dominated by deformed basaltic extrusive and shallowintrusive rocks intercalated with abundant chemical sedimentsincluding quartz–magnetite banded iron formation (BIF) and rareclastic sediments. The lithologies with ambiguous affinities aredominated by ultramafic and felsic rocks. The ultramafic rocks comein a great variety of types depending on metamorphic mineralogy, butthey can be divided into metadunites and metaperidotites based ontheir geochemistry and field appearance. They may represent mantle-derived tectonic slivers, ultramafic sills, and cumulates from maficintrusions, although it has not been possible to assign any specificaffinity to most individual bodies in the field. The felsic rockscommonly share petrographic and geochemical similarities with theAmitsoq gneiss that encloses the supracrustal belt. They includedetrital sediments and volcaniclastic flows but are dominated bygranitoid intrusions into the supracrustal pile and slivers of theAmitsoq gneiss that were tectonically intercalated with the supra-crustal rocks during late-stage orogenic events. The third categoryrepresents metasomatic rocks, which are derived from all lithologiesof the former categories by reaction with hydrothermal solutionsduring one or more tectonometamorphic events (Rose et al., 1996;Rosing et al., 1996). The metasomatic transformations have beenlocally extensive, and some lithologies that have been characterized asmetasediments in the literature can be shown to be metasomatic inorigin (Rosing et al., 1996; Myers, 2001). Some of these metasomatitesinclude metacarbonate rocks, garnet–biotite schists, and muscoviteand fuchsite schists.

A prominent unit in the ISB is the Garbenschiefer amphibolites(GA) (Fig. 1). This unit is composed of schistose metabasic rocks thatdisplay distinctive amphibole sheave textures (defined by bundles ofamphibole crystals set in a matrix of plagioclase and quartz) in mostoutcrops. Chlorite is present in many Garbenschiefer rocks. This hasled to the common suggestion that the Garbenschiefer Unit representsa greenschist facies domain (Hayashi et al., 2000; Komiya et al., 2002;Appel et al., 2003) within the otherwise dominantly amphibolitefacies of the ISB (Boak and Dymek, 1982; Rollinson et al., 2002).However, the chlorites in the Garbenschiefer Unit are Mg-rich,reflecting the high Mg/Mg+Fe and high Al content of the amphibo-lites. Magnesian chlorite is stable well into amphibolite facies(Winkler, 1967), and there is no basis for separating the Garbenschie-fer Unit out as a separate metamorphic zone. The distinctivemetamorphic paragenesis discriminates the Garbenschiefer metaba-sic rocks from “normal” hornblende–plagioclase–quartz amphibolitesthat are common throughout the belt. These metabasites are derivedfrom tholeiitic extrusive and shallow intrusive protoliths (Gill andBridgwater, 1979; Nutman et al., 1984; Gill et al., 1988; Komiya et al.,2004). The tholeiitic metabasic rocks are everywhere stronglydeformed, but in places relic pillow structures (often with ocelli)and dike-in-dike relationships are preserved. In the SWpart of the beltpillow lava and dike lithologies are followed along-strike by a largearea of homogeneous meta-gabbroic amphibolite. These rock

sequences were interpreted as the deformed and disrupted fragmentsof a Paleoarchaean ophiolite complex (Furnes et al., 2007a). Thesedeformed rock suites are in turn intruded by later undeformedAmeralik dikes (Nutman and Friend, 2007; Furnes et al., 2007b).

2.2. Age relations and deformation

The Isua supracrustal belt has been deformed andmetamorphosedduring several distinct events. Penetrative deformation and amphibo-lite facies metamorphism affected all supracrustal units and theenveloping gneiss complex prior to the emplacement of the basalticAmeralik dike swarm ca. 3550 Ma ago (Gill and Bridgwater, 1979;Nutman et al., 2004). These deformed amphibolites are easily dis-tinguished from later undeformed dikes, such as the ubiquitous setof basaltic, plagioclase-phyric Ameralik dikes, both petrographicallyand in the field (Furnes et al., 2007b). This pre-Ameralik tectono-metamorphic event was contemporaneous with a major phase ofAmitsoq gneiss magmatism ca. 3750 Ma (Crowley et al., 2002; Freiet al., 2002; Nutman et al., 2002, 2007), which likely provided the heatfor metamorphism and the production of metasomatic fluids. Thedeformation, metasomatism, and metamorphism associated with thisevent have obliterated most primary contacts between the supracrus-tal belt and the enveloping gneisses, and between different lithostrati-graphic units within the belt. However, the precursors of the Amitsoqgneiss were probably dominantly intrusive into the supracrustalbelt, as evidenced by several well-preserved intrusive contacts withgneisses ranging in age from 3810Ma to 3710Ma (Nutman et al., 1993,1997, 2007; Crowley, 2003).

Whether the ISB constitutes one continuous volcano-sedimentarysuccession or whether it comprises tectonic units with different agesand origins has not been well established in the literature. Isotopicages of the supracrustal lithologies are strongly influenced bymetamorphic and metasomatic overprinting, and while often precisethese data are unlikely to reflect accurate ages of different units.Consequently, the age of the ISB rocks is mainly constrained by U–Pbzircon ages of crosscutting felsic intrusives. This principle has in itselfbeen the focus of major controversy because of the possibility ofinheritance in the zircon populations of felsic rocks (Kamber andMoorbath, 1998; Nutman et al., 2000; Whitehouse et al., 2001).However, most felsic units studied have simple igneous zirconpopulations with tightly clustered age distributions. From theavailable data, there are a large number of felsic dikes with dates ofca. 3800 Ma in the ISB near the outward convex border of the belt,whereas the studied felsic dikes near the inward concave margin ofthe belt have dates close to 3700 Ma. This has given rise to the ideathat the belt is composed of at least two tectonostratigraphic unitsseparated by 100 Ma in time of formation (Nutman et al., 1997, 2007).The finding of 3690–3700 Ma zircons in BIF and felsic rocks in theinferred inner unit has been taken as evidence for the ca 3700 Ma ageof this part of the ISB. However, the significance of these zircon datesbuilds on the interpretation of the felsic rocks as extrusive and thezircons in the BIF as detrital. Both claims are contentious and have noindependent observational support. In the absence of more robustevidence for a composite origin of the ISB, we treat it as onetectonostratigraphic unit constrained to be older than the envelopingAmitsoq/Itsaq gneiss complex (Fig. 1) by intrusive relationships alongboth the inner and outer margins.

3. Field relationships of the mafic–ultramafic rocks of the ISB

For the field and geochemical description of magmatic rocks in theISB we basically maintain the subdivision of the complex into central,outer and inner arc domains, as defined by Polat et al. (2002) and Polatand Hofmann (2003). However, this description is not entirelyrepresentative when it comes to the eastern arm of ISB, and wetherefore propose the following terms for the units: 1.


Undifferentiated amphibolites (UA) for the formerly named “outerand inner domains”, and 2. Garbenschiefer amphibolites (GA) (Fig. 1)for the “central domain”. In the following we consequently use thesetwo terms, UA and GA.

3.1. Undifferentiated amphibolites (UA)

All our field data pertaining to our interpretation of the ISB as anophiolite come from the undifferentiated amphibolites of the westernarm of the ISB. The rocks in this region are exposed along the easternside of the belt referred to as the “inner arc tectonic domain” by Polatand Frei (2005). Below are brief field and petrographic descriptions ofthe lithological components (of the eastern part of the western arm)that together suggest the existence of an ophiolite—the Isua ophiolite(Furnes et al., 2007a).

3.1.1. Pillow lavaPillow structures can be found at many outcrops around locality 1

(Fig. 1). In general they are highly deformed, and a penetrativecleavage occurs parallel to the lithological layering of different rockunits. However, in places (at locality 1, see Fig. 1) the pillows arerelatively well-preserved, and between pillows, small pockets of inter-pillow hyaloclastite are common (Fig. 2A, B). Ocelli-bearing pillows(Fig. 2C), characteristic of Archean pillow lavas (e.g. de Wit andAshwal, 1997), are useful for estimating the degree of deformationthese rocks experienced. Deformed ocelli (originally sphericalobjects) (Fig. 2D) indicate deformation with 80–90% shortening

Fig. 2. A–C. Pillow lavas, interpillow hyaloclastite (IPH) and ocelli. A pillow with dense oceLocation: Western arm of Isua Supracrustal Belt, location 1. Scale information: A. Pocket kniC. GPS bag in lower left is ca. 12 cm long. D. Hammer head is 13 cm long.

perpendicular to the cleavage and 200–250% extension along a welldefined lineation plunging at 72 S.

3.1.2. DikesAt locality 2 (see Fig. 1) the sequence consists of tabular, sub-

parallel dikes with intervening cm- to dm-thick zones of lenticular toirregular screens of volcanic material (Fig. 3). Approximately 500 mfurther south, at locality 3 (see Fig. 1), the mixed dike/volcanicsequence changes structurally downward into a sheeted complex,which consists of 100% tabular dikes (Fig. 4A–C). This dike complex isin tectonic contact withmetagabbro and ultramafic sheets to thewest.Individual dikes range in width from 2 to 50 cm. Dikes have both one-and (mostly) two-sided, fine grained, chilled planar margins (Fig. 4C).Crosscutting dikes are also observed (Fig. 4D). This sequence wasinterpreted as part of a sheeted dike complex by Furnes et al. (2007a).

Petrographically the central parts of sheeted dikes consist of fine-grained (grain size ∼300 μm) plagioclase, amphibole (predominant),and biotite with remnant subophitic textures. Dark green, commonlyschistose, marginal zones are interpreted to represent chilled margins(Fig. 4C) and are composed of dense (∼100 μm)monomineralic zonesof amphibole. These chilled margins of dikes are mineralogically andtexturally similar to the chilled margins of pillows in this area.

3.1.3. PlagiogranitePlagiogranite occurs as small pockets associated with the sheeted

dike complex (Fig. 5A). Petrographically it contains phenocrysts ofplagoioclase and amphibole set in a fine-grained quartz–plagioclase

lli-development in the central part is shown in C. D. Deformed ocelli within a pillow.fe in upper left corner is 9 cm long. B. Number label in lower right corner is 10 cm long.

Fig. 3. Dikes and screens of volcanic rocks at location 2 in Fig. 1. Length of hammer shaft is ca. 60 cm.


matrix (Fig. 5B, C). The geochemical composition of this rock type(sampled from the plagiogranite occurrence shown in Fig. 5A) is givenin Furnes et al. (2007a).

Fig. 4. Dike complex of the Isua Supracrustal belt. A. Ca. 30 m long continuous outcrop sectiondike units. C. Detail of individual dikes with pronounced chilled margins. D. Cross-cutting d

3.1.4. GabbroMetagabbro is rather rare in the ISB. However, in the southernmost

part of thewestern arm of the ISB, a relatively small area of metagabbro

(in the foreground of the picture) across part of the sheeted dike complex. B. A series ofikes. From location 3 on Fig. 1. Hammer head in B, C and D is 13 cm long.

Fig. 5. A. Plagiogranite (pale grey, irregular pods) associated with dikes and screens of volcanic rocks. Hammer for scale (32 cm long shaft). From location 2 on Fig. 1. B and C.Photomicrographs showing the quartz–albite–amphibole groundmass with plagioclase phenocrysts (Pl) in B, and an amphibole phenocryst (Amph) in C.


is exposed (Fig. 6). Locally this gabbro displays distinctive layering(Fig. 6B) and has been transformed to a flasergabbro due to solid-statedeformation as evidenced by the elongated streaks of plagioclase(Fig. 6C).

3.1.5. Ultramafic rocksUltramafic rocks constitute a significant part of both the western

and eastern arms of the ISB (Fig. 1). They are composed of medium tocoarse grained, red to dark grey/black rocks with alternating reddishand dark grey bands of various thicknesses (Fig. 7). In general theunspecified ultramafic rocks are enveloped by calc-silicate rocks. Thecalc-silicates are interpreted to have been formed by carbonationand desilication of country rock by fluids flowing out of the ultra-mafic rocks (Rose et al., 1996; Rosing et al., 1996). Possible protolithsof the highly altered ultramafic rocks (mantle or cumulates) arehence difficult to decipher, although locally they may representultramafic sills within the volcanosedimentary units (Dilek andPolat, 2008).

3.2. Garbenschiefer amphibolites (GA)

The Garbenschiefer amphibolites (GA) define a significant part ofthe ISB (Fig. 1). Mostly this unit consists entirely of recrystallizedrocks, mostly highly schistose, composed of a hornblende–garnet–biotite–chlorite mineral assemblage (Fig. 8A, B) (Rosing, 1999).Although the protoliths of these rocks are hard to recognize, it isstill possible to distinguish locally between metavolcanic/intrusiveand metasedimentary (volcaniclastic) protoliths. In the lowest-straindomains well-preserved sedimentary structures such as graded-bedding can be found (Fig. 8C, D) (Rosing,1999), and pillow structureshave been reported (Komiya and Maruyama, 1995).

4. Geochemistry

In this section, we summarise the published, available geochemicaldata from themeta-ultramafic tometabasaltic rocks of the GA and UAofthe ISB, togetherwithnewdata from theGA (Table 1). The sampleswerefused with LiBO2 and dissolved in HNO3. Major and minor elementswere analysed by ICP-Emission and all the trace elements by ICP-MS atSARM (Service d'Analyses Roches et des Mineraux) at Centre deResearches Petrographic et Geochemique, Vendoeuvre, France. Wealso compare these ISB data with representative geochemical data fromother ophiolite types.

4.1. Isua geochemistry

Fig. 9 shows Bowen diagrams for somemajor (SiO2, TiO2, Al2O3, FeOt,CaO) and trace (Cr, Ni, Zr) elements. Both with respect to the major andtraceelements themagmatic rocksdefine twodistinct trends. This iswelldemonstrated with respect to TiO2, Al2O3, FeOt and to some extent CaO,andZr. The rocksof theGA, togetherwith the rocks referred to as “MORB”by Komiya et al. (2004), have distinctly lower TiO2 and Zr, and generallylower FeOt and CaO, and higher Al2O3 contents than the UA of the ISB(outer and inner arc domains of Polat and Frei (2005)), and the “OIB”rocks of Komiya et al. (2004). Both types show continuous ranges in theMgO contents between ca. 17.5–5.5 wt.% for the GA rocks, and ca. 20.5–1 wt.% for the UA (Fig. 9). All the elements presented show relativelywell-defined magmatic trends with steadily increasing SiO2, TiO2, Al2O3,Zr, and decreasing Cr and Ni contents with decreasing MgO. WithdecreasingMgO, CaO increases up to ca. 7wt.%, and decreases thereafter.

Fig. 10 shows MORB-normalised multi-element diagrams of allavailable samples from the UA and GA of the ISB. The two former showratherflat, but slightly enrichedpatterns towards themost-incompatible

Fig. 6. Deformed metagabbro from the southernmost part of the western arm of ISB. A. General view of gabbro exposure. B. Rhythmically phase-layered gabbro. C. Close-up offlasergabbro, showing the lenticular appearance of plagioclase. Location of C is shown by the boxed area in A.


elements (to the left in the diagram), and also display minor negativeNb-anomalies and positive Pb-anomalies. The samples of the GA showhighly depleted and convex-downward MORB-normalized patterns,with slight negative Ta and Nb anomalies, and considerable positive Pbanomalies. Themost incompatible elements, i.e. Th and Ba,mostly showpronounced enrichment relative to the most depleted elements (e.g. Tithrough La of Fig. 10).

Fig. 11 shows themafic and ultramafic rocks in the tectonic domainsplotted inTh/Ybvs.Nb/Ybdiagram. In this diagramthe samples fromtheGA show a pronounced spread, but define a different field than thesamples from theUA. Both groups, however, plot above themantle arraydefined by the fields for N-MORB, E-MORB and OIB (Fig. 11).

4.2. Isua geochemistry compared to Phanerozoic ophiolites

In order to compare the ISB lithologies and the geochemicalcharacter of the pillow lavas and dikes with ophiolites, it is pertinent

to briefly review the most-recent classification of ophiolites based onPhanerozoic examples (Dilek, 2003). Subsequent to the definition ofthe Penrose-type ophiolite (Anonymous, 1972), a wealth of informa-tion on structural architecture, geochemical fingerprints, and evolu-tionary paths, that suggest different tectonic environments of ophioliteformation. Dilek (2003) compiled this information and proposedthe following classification into seven ophiolite types: (1) Ligurian-,(2) Chilean-, (3)Macquarie-, (4) Caribbean-, (5) Franciscan-, (6) Sierran-,and (7) Mediterranean (or suprasubduction zone)-type ophiolites.

As indicated above, Bowen diagrams of the ISB data (Fig. 9) arecompared with the data from two of the above-mentioned ophiolitetypes, i.e. the LIP-type Caribbean- and suprasubduction zone (SSZ)type Mediterranean ophiolites (Fig. 12). The UA of the ISB, particularlythe high-MgO samples, can be closely matched with volcanic rocksfrom the Cretaceous Caribbean ophiolites with Large Igneous Province(LIP) affinities. In the same diagram (Fig. 12) the ISB data are com-pared with the volcanic rocks from the Tethyan SSZ-type ophiolites.

Fig. 7.Ultramafic rocks from the eastern part of thewestern arm of the ISB (in the area between locations 1 and 3 in Fig.1). A. Large-scale, alternating reddish-brown (meta-dunitic to-harzburgitic) and black (metapyroxenite-rich) layers. B. Close-up photograph of the reddish-brown metaperidotite. C. Small-scale, alternating red and dark grey layers ofmetaperidotite. Scale information: A. Hammer shaft is ca. 60 cm long. B. Shoes in upper left are ca. 35 cm long. C. Black GPS bag is 12 cm long.


The large variations in the magmatic rocks from these ophiolite typesdefine broad fields with respect tomost elements. The ISB rocks do notoccupy the entire fields defined by these ophiolites in the Bowendiagrams, particularly the highly fractionated (SiO2-rich) rocks,although some important similarities appear. For example, thecomplete range in MgO and the low concentrations of TiO2 and Zr,and partly the high Al2O3 content of the rocks of the GA are highlysimilar to the trends defined by the typical boninitic to IAT rocks ofthese Tethyan SSZ-type ophiolites. It should be noted, however, that

none of the ophiolitic rocks match the high-Al2O3 rocks of the ISB inthe ca.14–16 wt.% MgO range.

Fig. 13 shows the various ophiolite types plotted in the Th/Yb vs.Nb/Yb diagram. Whereas the Ligurian- and Caribbean-types, andCalifornian ophiolites plot within the mantle array from N-MORB to E-MORB to OIB, the SSZ-type plot above the mantle array (Fig. 13).

With respect to the multi-element diagrams we have comparedthe UA and GA with the characteristic multi-element diagrams forCaribbean- and SSZ-type ophiolites (Fig. 14). The high-MgO basaltic

Fig. 8.Garbenschiefer amphibolites from thewestern arm of ISB. A. General viewof the Garbenschiefer amphibolite; B. Close-up from picture A, showing cm-long amphibole sheaves;C. Volcaniclastic sediments showing graded bedding; D. Close-up from picture C. Finger points to the grey graphite-bearing pelitic top of a graded bed (younging to the left). Allphotos from the western arm of ISB.


rocks from the Caribbean-type ophiolites compare well with those ofthe UA (Fig. 14). One difference between the Caribbean and the UA isthat the latter shows a negative Nb anomaly, a feature that is notrepresented by the Caribbean-type. The samples from the high-MgOrocks of the GA, on the other hand, show a pattern in the multi-element diagram that is nearly indistinguishable from a typical SSZ-type boninite (Fig. 14).

5. Geochemical and petrogenetic implications

5.1. Element mobility

During alteration and metamorphism of basaltic rocks elements arevariably mobilized (e.g. Cann, 1970; Coish, 1977; Humphris andThompson,1978). Several studies investigating elementmobility duringalteration andmetamorphism have identified, however, some elementsthat retain approximately their original concentration (Nicollet andAndribololona, 1980; Weaver and Tarney, 1981), or element ratios that

are little affected by metamorphism. Previous geochemical studies ofthe Isua ultramafic andmaficmagmatic rocks suggest that Zr, REE, Th, Ti,Nb and Yare relatively immobile (Polat et al., 2002; Polat and Hofmann,2003; Polat and Frei, 2005). This is consistent with other studies ofArchean volcanic rocks that also find that the relativemobilities of Al, Crand Ni are notably less than the commonly large scale mobility of Rb, K,Na, Sr, Ba, Si, Ca, Mg, Fe, P and Pb (see Polat et al., 2002 and referencestherein). The Bowen diagrams shown in Fig. 9 should, therefore, beinterpreted with care, since variations in Si, Mg, Fe and Ca may partlyhave resulted from secondary element mobility. However, the largerange in the concentrations of both major and trace elements as shownin Fig. 9 (Si, Ti, Al,Mg, Ca, Zr, Cr, Ni), and the relatively high stability of Al,Ti, Zr, Cr and Ni can hardly be accounted for by alteration processesalone. Moreover, all of the Bowen diagrams shown in Fig. 9, with theexception of the FeOt vs. MgO diagram, show systematic, positive orinverse correlationswithMgOcontent that are reminiscentof the trendscaused by fractional crystallization. Furthermore, the patterns definedby the multi-element diagrams (Fig. 10) and the Th/Yb vs. Nb/Yb

Table 1Major and trace element analyses of Garbenschiefer amphibolites from the Isua supracrustal complex.

Sample # 810381a 242790A 242670 242744 242743 242742 242733 242737 242729 242727H 242725 242719c 242718 242717A 242689 242673b 242671b

Si02 48.61 49.60 49.50 50.03 46.72 47.10 48.62 45.14 46.62 43.68 44.88 46.50 47.22 46.29 48.74 49.31 49.47Al203 17.73 17.60 18.52 17.82 16.34 19.56 17.37 15.99 13.42 9.91 14.52 16.83 16.08 15.11 17.83 16.19 18.20Fe203t 10.41 11.29 10.72 11.54 9.69 8.71 10.92 9.80 11.02 9.09 11.38 9.12 9.97 9.89 11.03 11.08 11.13MnO 0.13 0.17 0.23 0.17 0.14 0.13 0.15 0.15 0.23 0.15 0.23 0.20 0.16 0.19 0.22 0.19 0.19MgO 12.76 9.72 8.80 8.05 14.83 10.88 9.31 16.37 14.28 23.44 14.80 11.31 13.25 12.85 9.37 17.15 9.58CaO 7.57 8.10 9.32 7.37 6.28 8.07 8.60 5.84 10.17 7.22 8.71 11.68 7.13 9.74 9.33 2.78 9.02Na2O 0.96 2.00 1.23 1.80 0.66 1.68 1.65 1.13 0.62 0.07 0.95 1.18 1.02 0.78 1.88 0.50 1.43K20 0.15 0.02 0.08 0.15 0.02 0.10 0.10 0.04 0.07 0.15 0.04 0.06 0.05 0.10 0.11 0.02TiO2 0.27 0.32 0.31 0.39 0.21 0.21 0.36 0.20 0.19 0.14 0.21 0.23 0.22 0.16 0.30 0.22 0.29P205 0.18 0.21 0.21 0.20 0.19 0.20 0.22 0.18 0.20 0.17 0.22 0.23 0.21 0.18 0.19 0.15 0.18LOI 1.15 0.91 1.01 2.42 4.85 3.29 2.64 5.10 3.11 6.07 3.69 2.32 4.62 3.75 0.95 2.24 0.40Total 99.92 99.94 99.93 99.94 99.93 99.93 99.94 99.94 99.93 99.94 99.74 99.64 99.94 98.99 99.94 99.92 99.91Be 0.53 0.12 0.19 0.1 0.13 0.19 0.1 0.24 0.36V 207 217 209 233 185 198 230 160 157 136 184 183 177 163 238 184 211Cr 255 217 254 55.6 701 359 106 1207 1705 2203 1343 908 878 728 363 1165 271Co 68.0 67.9 62.7 45.4 63.4 44.8 53.7 65.8 73.8 81.8 77.7 61.3 70.3 56.0 65.9 73.7 62.3Ni 210 112 177 59.1 342 149 122 472 625 811 586 366 403 330 174 469 160Cu 52.1 1.1 2.6 4.3 2.5 2.7 47.2 11.9 23.8 20.1 0.8 18.5 18.0 14.9 24.6 83.9 1.8Zn 64.8 81.8 77.7 78.0 62.2 52.1 49.5 59.3 101 75.1 84.1 60.6 66.4 56.7 84.3 74.9 65.7Ga 12.9 14.6 15.1 14.1 11.0 12.3 14.4 10.0 9.84 7.61 10.9 11.1 10.8 9.43 14.6 12.1 14.0Ge 1.66 1.65 1.54 1.43 1.65 1.97 1.91 1.35 1.76 2.22 1.63 1.26 1.32 1.33 1.69 1.55 1.43As 0.58 0.47 0.63 0.36 0.27 0.06 0.87 0.31 0.63 5.0 0.48 0.9 0.51 0.96 0.63 0.7 0.48Rb 10.01 0.71 4.87 3.28 1.08 1.24 0.97 1.51 1.49 0.49 4.36 1.15 2.38 1.37 6.78 16.29 1.50Sr 48.4 65.9 74.3 84.1 52.2 72.8 56.8 54.7 25.3 1.32 36.5 115 58.8 60.0 52.5 8.45 53.8Y 11.1 12.5 13.3 15.6 10.2 9.54 13.3 7.42 7.38 6.28 7.93 10.1 10.0 9.1 12.9 11.0 13.6Zr 14.5 20.3 25.7 23.4 12.5 15.3 19.3 10.4 13.1 6.25 9.82 12.9 12.8 10.4 21.2 12.5 24.0Nb 0.43 0.45 0.79 0.55 0.23 0.39 0.39 0.21 0.28 0.13 0.14 0.25 0.26 0.16 0.40 0.22 0.64Mo 0.13 0.14 0.22 0.21 0.11 0.09 0.13 0.21 0.25 0.22 0.15 0.10 0.20 0.17 0.27 0.30 0.24Cd 0.11 0.01 0.05 0.04 0.08 0.13 0.04 0.15 0.11 0.07 0.10 0.06In 0.06 0.08 0.06 0.11 0.07 0.08 0.09 0.09 0.07 0.06 0.09 0.07 0.06 0.05 0.07 0.05 0.05Sn 0.20 0.36 0.32 0.44 0.26 0.35 0.18 0.51 0.25 0.18 0.32 0.09 0.09 0.13 0.57 0.17 0.44Sb 0.17 0.29 0.26 0.13 0.20 0.05 0.21 0.22 0.31 0.40 0.30 0.44 0.27 0.22 0.31 0.09 0.23Cs 0.923 0.090 0.827 0.163 0.096 0.108 0.094 0.116 0.140 0.078 0.166 0.108 0.157 0.749 0.299Ba 8.4 12 11 22 5.0 10 4.8 3.4 6.8 0.6 29 9.8 15 7.5 1.5 4.1 3.7Hf 0.43 0.51 0.67 0.75 0.35 0.46 0.54 0.33 0.43 0.22 0.3 0.37 0.4 0.32 0.65 0.4 0.68Ta 0.0718 0.0468 0.0647 0.0480 0.0179 0.0318 0.0315 0.0236 0.0247 0.0101 0.0147 0.0241 0.0314 0.0175 0.0487 0.0208 0.0602W 66.3 107 75.5 0.22 29.8 0.05 0.06 30.4 0.11 16.2 38.6 90 88.0 36.5 69.0 42.0 82.8Pb 85.6 6.29 3.36 1.22 0.53 1.05 1.0 0.63 4.95 0.98 1.73 3.92 2.57 2.82 3.86 1.79 2.67Bi 0.05 0.07 0.01 0.0 0.0 0.0 0.02 0.01 0.02 0.03 0.03 0.02 0.52 0.13 0.02Th 0.140 0.140 0.430 0.180 0.022 0.170 0.043 0.067 0.010 0.022 0.052 0.063 0.053 0.230 0.016 0.260U 0.046 0.030 0.110 0.014 0.011 0.008 0.003 0.043 0.052 0.002 0.022 0.120 0.058La 0.85 0.96 2.08 1.25 0.59 1.02 0.57 0.13 0.40 0.35 0.32 0.37 0.61 0.51 1.13 0.33 1.90Ce 1.82 2.51 4.53 3.21 1.22 2.47 1.88 0.38 0.89 0.27 0.96 1.03 1.64 1.09 3.32 1.09 4.24Pr 0.23 0.38 0.59 0.47 0.20 0.31 0.33 0.09 0.14 0.11 0.18 0.20 0.23 0.21 0.43 0.19 0.56Nd 1.15 1.73 2.62 2.31 0.95 1.35 1.91 0.45 0.55 0.62 0.83 1.07 1.24 0.82 1.96 1.17 2.75Sm 0.55 0.69 0.82 0.86 0.37 0.54 0.64 0.26 0.19 0.21 0.38 0.41 0.43 0.36 0.69 0.51 0.86Eu 0.16 0.21 0.29 0.21 0.16 0.17 0.22 0.10 0.08 0.04 0.11 0.15 0.15 0.12 0.27 0.16 0.29Gd 0.73 0.98 1.07 1.15 0.59 0.64 1.05 0.48 0.40 0.38 0.59 0.74 0.70 0.61 1.13 0.83 1.16Tb 0.16 0.20 0.22 0.22 0.15 0.17 0.22 0.09 0.08 0.09 0.11 0.14 0.15 0.14 0.21 0.15 0.23Dy 1.35 1.56 1.54 2.06 1.16 1.27 1.56 0.81 0.75 0.64 0.98 1.20 1.25 1.10 1.66 1.37 1.68Ho 0.37 0.40 0.43 0.49 0.32 0.27 0.41 0.23 0.22 0.18 0.28 0.31 0.33 0.26 0.40 0.32 0.45Er 1.19 1.24 1.29 1.53 1.00 0.99 1.28 0.83 0.78 0.61 0.83 0.97 1.08 0.83 1.21 1.13 1.36Tm 0.20 0.22 0.25 0.27 0.16 0.18 0.23 0.14 0.13 0.11 0.17 0.15 0.16 0.18 0.19 0.17 0.23Yb 1.37 1.32 1.65 1.91 1.29 1.14 1.36 0.92 1.08 0.84 0.98 1.20 1.28 1.11 1.58 1.33 1.54Lu 0.23 0.23 0.27 0.27 0.23 0.19 0.22 0.16 0.18 0.16 0.14 0.21 0.18 0.18 0.26 0.22 0.26

Fe2O3t = total iron as Fe2O3. LOI = loss on ignition.

a Sample 810381 is from the eastern arm of the ISB; all the others are from the western arm of ISB.

124H.Furnes

etal./

Lithos113

(2009)115

–132

Fig. 9. Bowen diagrams for the mafic and ultramafic magmatic rocks of the Isua supracrustal belt (ISB). Data are from: Polat et al. (2002); Polat and Hofmann (2003); Komiya et al.(2004); Furnes et al. (2007a); this work (Table 1). The undifferentiated amphibolites (1) and (2) are for the western and eastern part, respectively, of the western arm of the ISB(Polat and Hofmann, 2003).


diagram (Fig. 11) are based predominantly on elements that areconsidered to be stable during metamorphism (except Ba and Pb).However, Rosing and Frei (2004) showed extensive metasomatic Thenrichment in some Isua rocks. Based on 208/204–206/204 Pb isotopesystematics, Frei et al. (2002) demonstrated that Th was mobilized andintroduced into these rocks during 2800 Ma tectonometamorphicevents that affected most Isua rocks. It may, therefore, be expected that

Fig.10.Multi-element diagramsof representative samples fromtheUndifferentiatedamphibolitHoffman (2003); C: Polat et al. (2002) and this work (Table 1). NormalizingMORB values (in ppPb (0.3), Pr (1,32); Sr (90), Nd (7.3), Zr (74), Sm (2.63), Eu (1.02), Ti (7620), Gd (3.68), Tb (0.67

some of the samples analyzed in this study have experienced Th-enrichment above their primary values.

5.2. The ophiolite rock association

The lithological components of the UA, namely the pillow lavas,sheeted dikes, transitional zone between volcanic rocks and dikes,

es (UA) (AandB), andGarbenschiefer amphibolites (GA) (C). Data from:A andB:Polat andm) (after Pearce and Parkinson,1993) are: Ba (6,3); Th (0.12), Nb (2.33), La (2.5), Ce (7.5),), Dy (4.55), Y (28), Ho (1.01), Er (2.97), Tm (0.456), Yb (3.05), V (300), Cr (275), Ni (100).

Fig.11. The Isua data plotted in Th/Yb vs. Nb/Yb diagram.ModernMORB (N-MORB and E-MORB) and OIB define a diagonal arraywith N-MORB, E-MORB and OIB at its centre. Magmasthat have interacted with continental crust on ascent, or have a subduction component, are, at a given Nb/Yb ratio, displaced to higher Th/Yb values. After Pearce (2008). Isua datafrom: Polat et al. (2002); Polat and Hofmann (2003); this work (Table 1). The undifferentiated amphibolites (1) and (2) are for the western and eastern part, respectively, of thewestern arm of the ISB (Polat and Hofmann, 2003).The field defined by the Mariana arc is taken from Pearce (2008).


plagiogranite, gabbro and ultramafic rocks (Figs. 2–7), togethercomprise all the components of a complete Penrose-type ophiolite(Dilek, 2003). Furnes et al. (2007a) suggested that this rockassociation, together with geochemical affinities of pillow lavas anddikes, represents an ophiolite. Given that this rock sequence has beenstrongly deformed it is admittedly difficult on the basis of fieldrelationships alone to prove that these rock components once were acoherent slab of oceanic crust. Current geochronological constraintson the UA do not however, exclude the cogenetic origin of the variouslithological components. Furthermore, the geochemistry of themetabasaltic rocks of the UA (Fig. 9), shows a continuous composi-

Fig. 12. Bowen diagrams for different types of Phanerozoic ophiolites (suprasubduction zGarbenschiefer amphibolites (GA). Data sources for the Isua amphibolites are provided in(2008), Pindos: Pe-Piper et al. (2004); Saccani and Photiades (2004), Oman: Lippard et al. (1Auclair and Ludden (1987); Taylor (1990), Kizildag: Dilek and Thy (1998); Y. Dilek (unpubl

tional range from high- to low-Mg basalts represented by dikes andpillow lavas, and thus strengthen the cogenetic origin of theselithological units. On the basis of the present knowledge of the UA, wethus maintain our previous conclusion that this unit represents adismembered Archean ophiolite (Furnes et al., 2007a, b).

5.3. Geochemical characteristics

The geochemical characteristics of themetabasaltic rocks of the UAare, in most respects, compatible with the magmatic evolution of theMORB-type Ligurian and LIP-type Caribbean ophiolites (Dilek, 2003;

one and Caribbean types), shown together with the Isua undifferentiated- (UA) andFig. 9. The suprasubduction zone type ophiolites are represented Mirdita: Dilek et al.986); Einaudi et al. (2003); Godard et al. (2003); Troodos: Rautenschlein et al. (1985);ished data). Caribbean-type data: Klaver (1987); Kerr et al. (1996).

Fig. 13. The Phanerozoic ophiolite data plotted in Th/Yb vs. Nb/Yb diagram, onto which the fields for the two Isua populations are shown. See Fig. 11 for further information on datasources for the Isua amphibolites. Data sources for ophiolites are: Suprasubduction zone and Caribbean types: see Fig. 12 Ligurian ophiolites: Ferrara et al. (1976); Beccaluva et al.(1977); Ottonello et al. (1984); Vannucci et al. (1993); Rampone et al. (1998); Californian and Philippine ophiolites: Harper (1984, 2003), Harper et al. (1988); Dilek et al. (1991);Evans et al. (1991); Yumul et al. (2000); Metzger et al. (2002); Shervais (1990); Giaramita et al. (1998); Shervais et al. (2005).


Figs. 12 and 13). The MORB-type Ligurian ophiolites represent theearly stages of opening of an oceanic basin, whereas the LIP-typeCaribbean ophiolites represent the oceanic crustal assemblages of an

Fig. 14. Trace element comparison between amphibolites of the Isua supracrustal belt annormalized multi-element diagram of the undifferentiated amphibolites (UA) compared wiGarbenschiefer amphibolites (GA) compared with a typical boninite sample from a suprassources for the Undifferentiated- and the Garbenschiefer amphibolites, see Fig. 10.

oceanic plateau (Dilek, 2003). None of these ophiolite types areassociated with subduction-related magmatism. However, in the Th/Yb vs. Nb/Yb diagram (Fig.11) all of the analyses plot above themantle

d representative basaltic and boninitic rocks from Phanerozoic ophiolites. A. MORB-th basaltic samples from the Caribbean type ophiolite. B. Multi-element diagram of theubduction zone type ophiolite (Mirdita ophiolite, Albania, see Dilek et al., 2008). Data


array, a feature that may result from crustal contamination orsubduction-related recycling influence. The same conclusion may beinferred from the small negative Nb-anomalies (Figs. 10 and 14).

The geochemical characteristics of the rocks of the GA showsimilarity with the boninitic to IAT magmatic rocks of SSZ-typeophiolites in theMediterranean region (Figs. 12–14). As in the case ofthe rocks of the UA, whether this geochemical trend was influencedby crustal contamination or subduction processes, is a fundamentalquestion. Although the commonly-accepted evidence for the oldestcontinental crust is around 4 Ga (Bowring andWilliams, 1999), morerecent Hf-isotope studies of ca. 4.4 Ga zircons from Jack Hills,Western Australia, indicate that continental crust may be traced backfor an additional 400 My (Harrison et al., 2005). Therefore, it isprobable that at the time of formation of the Isua rocks (UA and GA),continental crust might have been recycled into the mantle.However, the amount of recycled continental crust was probablyminor in the early stages of the Isuan time span (ca. 3800–3500 Ma),and the Nd-isotope studies of the metabasalts (see below) do notfavour significant continental involvement in their genesis. It is,therefore, considered more likely that the negative Nb-anomalies,the Th/Yb vs. Nb/Yb relationships, and typical boninite-like patternsof the GA originated as a result of geochemical processes in asubduction zone.

There are no systematic geochemical differences between theamphibolites from the UA, but there are marked differences in majorand trace element compositions between the UA, and the rocks of theGA. Interestingly, the pillow lava amphibolites from the UA are not onlyindistinguishable in major and trace element patterns but also share a

Fig. 15. Schematic model showing proposed tectonic environment for the generation oamphibolites) of magmatic rocks of the Isua supracrustal belt, and their geochemical andmodification of melts (in A and B) are similar to a model suggested by Grove et al. (1992).

globally unique 0.3 e-unit 142Nd/144Nd anomaly, and together theydefine a statistically significant 3.78±0.04 Ga Sm–Nd isochron (Boyetet al., 2003). We find these geochemical and isotopic similarities instrong support of a similar age and origin for UA and consider thisevidence stronger than those arguments for their dissimilarity based onthe apparent lack of observed ∼3800 Ma felsic dikes crosscutting theseamphibolites. Consequently, we regard theUA as part of a single basalticvolcanic complex formed at ca. 3800 Ma.

The 143Nd/144Nd systematics of the Isua metabasic rocks has been amatter of continued attention duringmore than 30years of geochemicalresearch. Themain conclusionwe derive from these studies is that therehas been one or more phases of severe metasomatic disturbance of theSm/Nd isotopic system, and that the metabasaltic rocks most likely hadan initial positive 143Nd/144Nd anomaly of ca. + 1.5 e-units (Hamiltonet al., 1978; Gruau et al., 1996). This has been taken as evidence for theexistence of a MORB-like depleted mantle reservoir at the time offormation of the Isua rocks (e.g. Hamilton et al., 1983).

6. Geodynamic model: discussion and conclusions

The term suprasubduction zone (SSZ) was first introduced by Pearceet al. (1984) for ophiolites that have the crustal components andarchitecture of oceanic crust with geochemical signatures of subductioneffects. These ophiolites have been interpreted to have formed in intra-oceanic immature arc systems. Hawkins and Evans (1983) and Hawkinset al. (1984) expanded themeaningof the termsuprasubduction zone toinclude all of the components formed in a subduction-related systemfrom arc, forearc to backarc. A good example of this SSZ environment is

f the two units (UA–the Undifferentiated amphibolites, and GA–the Garbenschiefergeological characteristics. Formation and subsequent aggregation, differentiation and


the West Philippine Basin–Mariana trench system, where subductionrollback over the last 48 My has generated a series of backarc basins,remnant arcs and a frontal arc system (e.g. Crawford et al., 1981; Fryer,1992). In this convergent margin tectonic setting, the arc–forearc–backarc systems have been undergoing lithospheric extension. Whenslab rollback exceeds the plate convergence rates, extensional condi-tions prevail, producing seafloor spreading generated oceanic crust(Leitch, 1984; Dilek et al., 2008).

We propose a new geodynamic model for the magmatic develop-ment of the UA and GA in the Isua supracrustal belt (Fig. 15). In thismodel the UA represents the oldest unit of the ISB, and the fieldgeology shows that this unit contains all the components of anophiolite. The amphibolites (pillow lavas and sheeted dikes) display aMORB-like geochemistry with a wide compositional range fromprimary to highly fractionatedmagmas.We suggest that the UAunit ofthe ISB originally represented a coherent slab of oceanic crust formedby seafloor spreading (Fig. 15A). Compared with Phanerozoicophiolites, the geochemistry of the UA unit is most comparable withthe Ligurian- and Caribbean-type ophiolites that represent oceaniccrust without subduction imprint. It is possible, however, that some ofthe UA magmas were influenced by subduction as indicated by minornegative Nb-anomalies (Fig. 10) and the relatively high Th/Yb ratios ata given Nb/Yb value (Fig. 11). In various discriminant diagrams (Ti–V,Zr–Zr/Y, Ti–Zr, Y–Cr), the UA pillow lavas and dikes straddle theboundary between mid-ocean ridge basalts (MORB), island arctholeiites (IAT) and boninites (Furnes et al., 2007a). To what extentthe UA crust was influenced by subduction processes depends on thelocation of the mantle source producing the melts relative to thesubduction zone. In our model we suggest that the oceanic crust of theUA–the Isua ophiolite–formed at an early stage and in a setting awayfrommajor influence of a subduction zone (Fig. 15A, and see Dilek andFlower, 2003). Hence, it was little affected by subduction processes.

The GA consists of volcaniclastic rocks, pillow lavas and gabbroicintrusions (collectively referred to as Garbenschiefer). The geochem-ical composition of the amphibolites defines a typical boninite-likepattern, with strong depletion in incompatible elements. It thusdiffers from the UA unit both with respect to the lithologicalcomponents, structural architecture, and the geochemistry. Boninitesare invariably associated with the early stages of intra-oceanic islandarc generation, and are generally attributed to extension of forearclithosphere in response to slab rollback (e.g. Crawford et al., 1989;Bedard et al., 1998). We suggest that the GA unit represents anextended incipient island-arc fed by the magmas that were generatedfrom a subduction-influenced, repeatedly depleted and hydrated,refractory mantle (Fig. 15B). This model thus suggests that the GAformed synchronously, or after the formation of the UA, reminiscent ofmost SSZ-type Tethyan ophiolites in the Mediterranean region (Dileket al., 2008). The geochemical composition and trends of the boninite-like amphibolites of the GA are similar to the late stage IAT andboninitic lavas and dikes of the SSZ ophiolites in the Mediterraneanregion (e.g. Dilek et al., 2008). These Phanerozoic ophiolites appear tohave formed from partial melting of relatively hot, hydrous andrepeatedly depleted, refractory peridotite in rapidly evolving supra-subduction zone mantle wedge (Dilek et al., 2008).

In our model melt generation, aggregation/mixing, and differen-tiation occurred at multiple levels below the crustal segments of UAand GA. The build-up of the crust formed by intrusions and extrusionsof magmas ranging from primary (and relatively primitive) todifferentiated compositions. This geochemical trend resulted frommelting of an ambient mantle that first produced MORB-like magmasof the UA, and subsequently from a depleted and hydrated,subduction-influenced mantle that produced the boninite-like mag-mas of the GA (Fig. 15). This proposed construction of the UA and GA,for which the lithological components of the two units and thegeochemical composition of their amphibolites are compatible, is inprinciple similar to the generation of Proterozoic and Phanerozoic

SSZ-type ophiolites and modern analogues (e.g. Hawkins, 2003;Pearce, 2003; Dilek et al., 2008). In this way the Isua ophiolite can bedefined as a composite segment of Archean oceanic crust derived froma first-stage generation of MORB-like magmas with little to moderatesubduction influence (UA), followed by the generation of IAT-boninitemagmas (GA) from a hydrated and depleted mantle.

Polat and Frei (2005) also proposed a suprasubduction geody-namic setting for the boninite-like and non-boninitic rocks of Isua. Intheir model, oceanic ridge subduction is proposed as a necessity togenerate sufficient heat for the production of the high-MgOmagmaticrocks. However, if the high-MgO magmatic rocks of the GA representtrue boninites generated above a hydrated mantle wedge, in a similarmanner to the production of modern boninites, the necessity forgenerating extra heat by oceanic ridge subduction for their generationmay not be required (e.g. Stone et al., 1997; Parman et al., 2001, 2003;Parman and Grove, 2004; Grove et al., 2006).

The inferred SSZ origin of the Isua supracrustal units has significantimplications for the ongoing debate about the timing of the onset ofthe modern plate-tectonic processes (seafloor spreading and subduc-tion) in the Archean. These rocks represent some of the oldest intactrocks on Earth, and their lithological components and mutualrelationships, combined with the geochemical signatures of thelavas and dikes, can be found in Phanerozoic ophiolite terrains. Thisobservation implies in turn that the magmatic and tectonic processesand geodynamic setting during the formation of the 3.8 Ga Isua beltmay have been similar to those of the Phanerozoic SSZ ophiolites.Therefore, we maintain that Phanerozoic-like plate tectonic wasoperative at 3.8 Ga (Furnes et al., 2007a,b; Dilek and Polat, 2008).

Acknowledgements

This study was financed by a series of research grants from theNorwegian Research Council, the Geological Museum of Copenhagen,the GFZ-Potsdam, and the Agouron foundation. We thank NicolaMcLoughlin for comments on an early version of the paper, and JaneEllingsen for helping with the illustrations. We further thank the tworeferees, Simon A. Wilde and Gouchun Zhao, and the guest editor ofthis issue, Paul T. Robinson, for constructive comments and sugges-tions that improved the manuscript. This is AEON contributionnumber 63.

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