Growth of granite–greenstone terranes at convergent margins, and stabilization of Archean cratons

ELSEVIER Tectonophysics 305 (1999) 43–73

Growth of granite–greenstone terranes at convergent margins,and stabilization of Archean cratons

Timothy M. Kusky a,Ł, Ali Polat b

a Boston University, Center for Remote Sensing, 725 Commonwealth Ave., Boston, MA 02215, USAb Department of Geological Sciences, University of Saskatchewan, now at Max-Plank Institut fur Chemie,

Postfach 3060, 55020 Mainz, Germany

Received 3 April 1998; accepted 27 July 1998

Abstract

Archean granite–greenstone terranes represent juvenile continental crust formed in a variety of plate tectonic settingsand metamorphosed through a complex series of structural and magmatic events. Most Archean granite greenstone terranesappear to have acquired their first-order structural and metamorphic characteristics at convergent plate margins, wherelarge accretionary wedges similar in aspect to the Chugach, Makran, and Altaids grew through offscraping and accretionof oceanic plateaux, oceanic crustal fragments, juvenile island arcs, rifted continental margins, and pelagic and terrigenoussediments. Buoyant slabs of parts of Archean oceanic lithosphere may have been underplated beneath these orogens,forming thick crustal roots characterized by interleaving between the depleted slabs and undepleted asthenosphere.Back-stepping of the subduction zones after accretion of plateaux and arcs caused the arcs magmatic fronts to migratetrenchward through the accretionary wedges. Dehydration of the subducting slabs hydrated the mantle wedges below thenew arcs and generated magmas (sanukitoid suite) in the mantle wedge, whereas other magmas (tonalite, trondhjemite,granodiorite or TTG suite) appear to have been generated by melting of hot young subducted slabs. Eventual collision ofthese juvenile orogens with other continental blocks formed anatectic granites, then thickened the crust beyond its ability tosupport its own mass, which initiated gravitational collapse and decompressional release of syn- to late-tectonic granitoidsfrom wedges of fertile mantle trapped between underplated oceanic lithospheric slabs, and aided in the cratonization ofthe granite–greenstone terranes. Deeply penetrating structural discontinuities such as shear zones and sutures providedpathways for fluids and granitoids to migrate into the mid- and upper-crust, forming ore deposits and plutons. Mostpreserved granite–greenstone terranes have been tectonically stable since the Archean, and form the cratonic interiors ofmany continents. 1999 Elsevier Science B.V. All rights reserved.

Keywords: Archean; greenstone belt; granites; tectosphere; cratons; ophiolite

1. Introduction

Understanding the origin of stable continentalcratons hinges upon recognizing which processes

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change the volume and composition of continentalcrust with time, and how and when juvenile crustevolved into stable continental crust. The evidencefrom the preserved record suggests that the pre-served continental landmass has been growing sincethe early Archean, although the relative rates andmechanisms of crustal recycling and crustal growth

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44 T.M. Kusky, A. Polat / Tectonophysics 305 (1999) 43–73

are not well known, and have been the focus of con-siderable geological debate (e.g., Fyfe, 1978; Deweyand Windley, 1981; Armstrong, 1991; McCullochand Bennett, 1994; Taylor and McLennan, 1995).The oldest rocks known on the planet are the circa4.0-Ga Acasta gneisses (Bowring et al., 1989, 1990)from the Anton terrane of the Slave Province (Kusky,1989). The Acasta gneisses are chemically evolved,and show trace and REE patterns similar to rocksformed in modern supra-subduction zone settings(Bowring et al., 1989, 1990). Maruyama et al. (1991)and Komiya et al. (1999) suggest that the oldestknown sedimentary sequence (the circa 3.8-Ga Isuasupracrustal assemblage) is an accretionary complex.A few circa 4.2-Ga zircon grains have been found

Fig. 1. World map showing distribution of Archean cratons and areas underlain by Precambrian crust (compiled from: Unesco, 1976;Condie, 1981; Hoffman, 1989, 1997; Abbott and Mencke, 1990; Goodwin, 1991, 1996; de Wit and Ashwal, 1997). Cratons labelled asfollows: 1 D Slave Province, 2 D Superior, 3 D Wyoming, 4 D Kaminak (Hearne), 5 D North Atlantic (Nain, Godthaab, Lewisian), 6 DGuiana, 7 D Central Brazil (Guapore), 8 D Atlantic (Sao Francisco), 9 D Ukranian, 10 D Baltic (Kola), 11 D Aldan, 12 D Anabar, 13D Sino–Korean, 14 D Tarim, 15 D Indian, 16 D Pilbara, 17 D Yilgarn, 18 D Gawler, 19 D Kaapvaal, 20 D Zimbabwe, 21 D Zambian,22 D Angolan, 23 D Kasai, 24 D Gabon, 25 D Kibalian, 26 D Uweinat, 27 D Liberian, 28 D Maritanian, 29 D Ouzzalian, 30 D NapierComplex, 31 D Prince Charles Mountains, 32 D Vestfold Hills, 33 D Heimefront Ranges, 34 D deeply buried Archean rocks of the EastEuropean Shield, 35 D Tajmyr.

(e.g., Froude et al., 1983), but it is not clear if thesewere ever parts of large continental land masses.Rudnick (1995) estimates that approximately half ofthe present mass of continental crust was extractedfrom the mantle during the Archean.

Exposed portions of the Archean lithosphere(Fig. 1) are broadly divisible into two main cat-egories (e.g., Goodwin, 1996). The first are the‘granite–greenstone’ terranes, containing variablydeformed assemblages of mafic volcanic=plutonicrocks, metasedimentary sequences, remnants ofolder quartzo-feldspathic gneissic rocks, and abun-dant late granitoids. The second main class of pre-served Archean lithosphere is found in the high-grade quartzo-feldspathic gneiss terranes. Relatively

T.M. Kusky, A. Polat / Tectonophysics 305 (1999) 43–73 45

little deformed and metamorphosed cratonic coversequences are found over and within both typesof Archean terrain, but are especially abundant onsouthern Africa’s Kaapvaal craton (Fig. 1, Brandland de Wit, 1997). Also included in this categoryare some thick and laterally extensive carbonateplatforms similar in aspect to Phanerozoic carbon-ate platforms (e.g., Grotzinger, 1989; Danielson andKaufman, 1996) indicating that parts of the Archeanlithosphere were stable, thermally subsiding plat-forms (Kusky and Hudleston, in press).

In this contribution we focus on describing thetectonic setting of juvenile Archean rocks fromArchean granite–greenstone terranes to understandmechanisms of continental growth, and we look atprocesses that converted these juvenile terranes intomature continental crust. Insight into these processeswill contribute to our understanding of why Archeangranite–greenstone terranes form the stable cratoniccores of many continental interiors (Fig. 1). Al-though we recognize the importance of processesincluding impact melting in the earliest Archean,sub-continental plume and hot-spot related magma-tism, these processes typically involve significantreworking of older crust with variable amounts ofjuvenile magmatic additions to the continental land-mass, and they probably involve relatively little newaddition of mass to the continents (but see McKenzieand Bickle, 1988 for an alternative view). The smallsize of Archean continents probably allowed mostplumes to be diverted to their edges (e.g., Gurnis,1988) where they would form oceanic plateaux thatare preserved abundantly in the geological record.Here, we examine convergent margin processes fromArchean granite–greenstone terranes, including ac-cretion of juvenile island arcs, oceanic plateaux,collapsed back-arc basins, oceanic crust, and migra-tion of arc=trench systems through accreted mag-matic and sedimentary sequences. We approach thissynthesis by emphasizing specific, well-documentedexamples of each type of process from different cra-tons, emphasizing the Superior, Slave, Yilgarn, andZimbabwe cratons. We also compare these processesto those of Proterozoic and Phanerozoic orogens toarrive at a general synthesis of the processes thatenable juvenile crust of granite–greenstone terranesto form the cratonic cores of so many continents.

2. Mechanisms of growth of juvenile crust atconvergent margins

2.1. Juvenile island arc accretion

Many Archean granite–greenstone terranes are in-terpreted as juvenile island arc sequences that grewabove subduction zones, and later amalgamated dur-ing collisional orogenesis to form new continentalcrust (Burke et al., 1976; de Wit and Ashwal, 1986,1997; Park, 1991; Kiyokawa and Taira, 1995). Theisland arc model for the origin of the continentalcrust is supported by geochemical studies, whichshow that the crust has a bulk composition similarto arcs (e.g., Taylor and McLennan, 1995; Rud-nick, 1995). As emphasized by Hamilton (1988,1998), island arcs are extremely complex systemsthat may exhibit episodes of distinctly different tec-tonics, including accretion of ophiolite fragments,oceanic plateaux, intra-arc extension with formationand preservation of back-arc and intra-arc basins(e.g., Saleeby, 1992; Aitchison and Flood, 1994;Tardy et al., 1994; Stern et al., 1995a,b; Vergara etal., 1995; Xue et al., 1996; O’Brien et al., 1997).Many juvenile arcs evolve into mature island arcsin which the magmatic front has migrated throughits own accretionary wedge, and many evolve intocontinental margin arcs after they collide with othercrustal fragments or continental nuclei.

The Schreiber–Hemlo greenstone belt is locatedin the northeastern section of the greenstone–granitoid Wawa subprovince of the SuperiorProvince, which extends from the Vermilion dis-trict of Minnesota in the west to the Kapuskasingstructural zone in the east (Fig. 2; Williams et al.,1991). The belt is composed of ca. 2750–2695 Mamafic to felsic volcanic, 2720–2690 Ma mafic tofelsic intrusive, and 2705–2697 Ma siliciclastic sed-imentary (turbidite) rocks (Fig. 3; Table 1; Corfuand Muir, 1989a; Corfu, 1989b; Williams et al.,1991; Fralick, 1997). Field relations indicate thatboth turbidites and volcanic sequences are fault-bounded (Polat et al., 1998). The Schreiber–Hemlogreenstone belt terrane is typical of many proposedArchean island arc terranes, and also shows evidencefor accretion of oceanic fragments, and migration ofthe magmatic front through the subduction=accretioncomplex. These processes are discussed in detail


Fig. 2. (a) Subprovince map of the Superior Province (modified after Card, 1990). (b) Simplified geological map of the study area,showing the location of the Schreiber–Hemlo belt (modified after Williams et al., 1991). LSHFZ D Lake Superior–Hemlo fault zone.Open rectangle shows the location of Fig. 2. TTG D tonalite–trondhjemite–granodiorite.

here, as an example of the complexities of juvenileisland arc accretion in the Archean, obscured by latersignificant geological events.

Volcanic rocks consist mainly of tholeiitic pillowbasalts, and 10-cm to 1.5-m-thick tholeiitic basalt

flows, komatiites and komatiitic basalts with wellpreserved spinifex texture, and tholeiitic and calc-alkaline basalts to rhyolites. There are minor cherts,iron formations, and gabbros within these mafic–ultramafic volcanic sequences. Chert layers, banded

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Fig. 3. Geological map of the Santoy Lake–Middleton area of the Schreiber–Hemlo greenstone belt. The northern part of the map is modified from Walker (1967). HW17 DHighway 17; CPR D Canadian Pacific Railroad.


Table 1Summary of the lithological, geochronological, geodynamic, and structural characteristics of the Schreiber–Hemlo greenstone belt (fromPolat et al., 1998)

Lithologicalunit(s)

Age Interpretedgeodynamicsetting

Origin Tectonostratigraphicnature

Deformationphase

Block–matrix relations(Ma)

Komatiites, Mg-to Fe-tholeiitesand tholeiiticgabbros

2750–2700 Oceanicplateau(s)

Mantle plumes Allochthonous D1,D2,D3 Exotic blocks of oceanplateau basalts in matrixof sheared turbidites.Native blocks of gabbrosin matrix of shearedbasalts

Siliciclasticturbidites

2705–2697 Convergentmargin

Island arc trench Allochthonous D1,D2,D3 Exotic(?) or nativeblocks of ryolites inmatrix of shearedturbidites

Tholeiitic tocalc-alkalinebasalts torhyolites

2720–2695 Oceanic islandarc

Slab to wedgemelting

Autochthonous toallochthonous

D1,D2,D3 Exotic blocks of ryolitesin matrix of shearedocean plateau basalts

TTG plutons 2720–2690 Magmatic arc Slab melting Autochthonous D2,D3 No block–matrixrelations observed

Mafic to felsicsills and dykes

2690–2680 Magmatic arc Slab to wedgemelting

Autochthonous D2,D3 Native blocks oftonalites and gabbros inmatrices of shearedturbidites and basalts

Lamprophyres 2680–2674 Accretion-relatedextension

Lithosphericmantle

Authochthonous D3,? No block–matrixrelations observed

iron formations and pillar basalts suggest a marineorigin for these volcanic sequences (see Isozaki etal., 1990; Thurston, 1990; Ohta et al., 1996). Thegeochemistry of tholeiitic basalts and associated ko-matiites suggests an intra-oceanic (oceanic plateau)geodynamic setting (Table 1; Polat et al., 1996a,1998). Other units of tholeiitic and calc-alkalinebasalts to rhyolites have geochemical characteristicsof intra-oceanic arcs (Table 1; Polat et al., 1998).Plateau and arc basalts cannot be distinguished inthe field. Both volcanic and sedimentary tectonicsequences are intruded by syn-kinematic, slab-derived high-Al, high La=Ybn tonalite–trondhjem-ite–granodiorite (TTGs; see Martin, 1986; Drum-mond and Defant, 1990; Kerrich et al., 1996), and

Fig. 4. Interpreted geological cross-sections corresponding to field traverses from the Steel Lake (A–A0) and Ripple Lake (B–B0) areas.Inset diagrams are based on detailed field observations. For locations see Fig. 3. For purposes of illustration the sizes of transposedfolds are exaggerated. Arc-derived distal turbidites and oceanic plateau fragments were tectonically juxtaposed and imbricated by D1

thrust faults. These thrust faults were overprinted by, and reactivated as, strike-slip faults during D2. Both sedimentary and volcanic unitsunderwent intense transposition, tectonic wedging, fragmentation, and mixing during D1 and D2, forming broken formations and tectonicmelanges.

subduction-derived gabbroic sills and dikes. Collec-tively these lithological sequences are cut by latekinematic, ca. 2680–2674 Ma lamprophyre dikes(Fig. 4; Table 1).

Sedimentary rocks consist of interbedded, dis-tal turbiditic greywacke sandstones, siltstones, andshales, with minor black chert layers (Walker, 1967;Purdon, 1995; Polat et al., 1998). The thicknessof turbidite beds ranges from a few centimeters toone meter. Sandstones tend to be more abundantand thicker than inter-bedded shales and siltstones.Conglomerates are minor, and occur primarily inturbidite channels. Primary sedimentary structures(e.g., parallel bedding, cross-bedding, grading) areintensely deformed throughout the belt (Figs. 3


and 4; Kuhns et al., 1994; Polat et al., 1998). Pet-rographic observations and high-precision ICP–MStrace element data suggest a magmatic arc prove-nance for these trench turbidites (Table 1; Polat etal., 1998).

At lower greenschist facies, the Schreiber sectionof the belt is less metamorphosed than the Hemloarea, for in the Hemlo area, amphibolite facies meta-morphism is predominant. Primary igneous texturesand pillow geometries of volcanic rocks are wellpreserved outside of major shear zones.

Overprinting relationships between various typesof structural fabrics (e.g., folds, shear zones, thrusts)suggest that the Schreiber–Hemlo greenstone beltunderwent three phases of deformation. Polat (1998)and Polat et al. (1998) provide detailed descriptionof these deformation phases. The earliest phase ofdeformation (D1) is defined by rotated thrust faults,SSE-verging, tight asymmetric folds, sheath folds,and steeply dipping foliation and associated mineralelongation lineations. D1 reflects tectonic imbrica-tion of oceanic plateaus, island arcs, and arc-derivedturbidites in a subduction–accretion complex. D2

is dominated by ENE-striking, right-lateral, orogen-parallel strike-slip faults, and steeply dipping thrustshear zones: the presence of coeval strike-slip andcompressional structures is consistent with a right-lateral transpressional deformation. Fragmentationof oceanic plateau, island arc, and trench turbiditedeposits resulted in broken formations and a tectonicmelange during D1 and D2 formation phases (Po-lat, 1998). Contacts between turbidite and volcanicbelts are typically marked by several meter-thickD2 strike-slip fault zones. Along these strike-slipfaults, D1 thrust fabrics such as S–C planar fabricsand asymmetric porphyroclasts are locally preserved,suggesting that the D1 thrust faults were overprintedby, and reactivated as, the D2 strike-slip faults(Figs. 3 and 4). The common orientation of S1 folia-tions and L1 lineations in the volcanic and sedimen-tary rocks, as well as the presence of the D1 thrustfaults and tight to isoclinal folds in both volcanicand sedimentary sequences suggest that turbiditesand mafic–ultramafic volcanic sequences were struc-turally juxtaposed and imbricated during D1.

The most pronounced D3 structure in the studyarea is the right-lateral Lake Superior–Hemlo faultzone (LSHFZ; Fig. 2; Hugon, 1984; Muir and El-

liott, 1987). D3 in the Schreiber–Hemlo region is at-tributed to subprovince accretion (Polat et al., 1998).

In Phanerozoic orogenic belts, the occurrenceof tectonically juxtaposed oceanic (e.g., mid-oceanridge basalts, seamounts, oceanic plateaus) and arc-derived trench turbidites is considered as evidencefor ancient subduction zones (see Isozaki et al., 1990;Polat et al., 1996b; Kusky et al., 1997a). Oceanicbasalts in Phanerozoic subduction–accretion com-plexes are often older than turbidites against whichthey are juxtaposed (Jones et al., 1993). Collec-tively, the occurrence of interleaved fragments ofoceanic plateau and arc-derived turbidites in the lateArchean Schreiber belt occurred by D1 convergentmargin tectonics and are comparable to Phanerozoicsubduction–accretion complexes of Japan (Isozakiet al., 1990; Jones et al., 1993; Kimura et al.,1994). Similarly, tectonically juxtaposed fragmentsof oceanic crust and arc-derived turbidites in vari-ous greenstone belts of the Superior Province havebeen interpreted as Archean subduction–accretioncomplexes (Hoffman, 1991; Kimura et al., 1993).

The late stages of D1 and D2 were accompaniedby the intrusion of syn-kinematic, mafic to felsicdikes and sills ranging from a few centimeters toseveral hundred meters in thickness, primarily alongthe D2 strike-slip faults and S1 planes. The ab-sence of D1 fabrics in these dikes and sills suggestthat their intrusion post-dates tectonic juxtapositionof trench turbidites with oceanic arc and plateaufragments. The geochemistry of the dikes and sillssuggests oceanic slab-derived melts for felsic intru-sions, and slab-dehydration, mantle wedge-derivedmelts for mafic and intermediate intrusions (Table 1;Kerrich et al., 1996; Polat et al., 1998). The in-trusion of subduction-derived, syn-kinematic maficto felsic igneous rocks into imbricated turbidite,arc, and oceanic plateau sequences suggests juxta-position of these contrasting lithologies above anArchean subduction zone. The transition from D1

to D2 reflects the evolution from thrust tectonicsabove a subduction zone to right-lateral transpres-sion. The intrusion of subduction-derived igneousrocks into subduction–accretion complexes is an im-portant feature of Phanerozoic subduction–accretioncomplexes, such as the Chugach accretionary com-plex of Alaska (Page et al., 1986; Bradley et al.,1994; Kusky et al., 1997a,b), the Shimanto accre-


tionary complex of Japan (Hasebe et al., 1993),the Altaid accretionary complexes of Central Asia(Sengor and Natal’in, 1996), and the Cordilleran ac-cretionary complexes of British Columbia (Hollisterand Andronicos, 1997).

The Schreiber–Hemlo greenstone belt evolved asa juvenile magmatic arc and forearc accretionarywedge above an Archean subduction zone. This arcsaw episodes of accretion of normal ‘MORB-like’oceanic crust, oceanic plateau basalt, trench tur-bidites, and was intruded by arc magmas when themagmatic front migrated through the forearc as anormal product of accretionary growth of the wedge.D1-related thrusts and=or D2 strike-slip faults in thefore-arc region of the Schreiber–Hemlo magmaticarc may have provided conduits for the uprising slaband wedge melts (TTG and sanukitoids, respectively;Wyllie et al., 1997; Rapp, 1997), and induced de-compressional partial melting in the sub-arc mantlewedge to produce the subarc mantle-derived gab-bros (Figs. 5 and 6). A close relationship betweenorogen-parallel strike-slip faulting and magmatismhas recently been recognized in several Phanerozoictranspressional orogenic belts, including the NorthAmerican Cordillera (Tikoff and Saint Blanquat,1997; Hollister and Andronicos, 1997; Kusky et al.,1997a), Japanese island arcs (Isozaki et al., 1990),and British Caledonides (Hutton and Reavy, 1992),suggesting that as in Phanerozoic counterparts, oro-gen-parallel strike-slip faulting in the Schreiber–Hemlo greenstone belt played an important role inlateral crustal accretion and magma emplacement.

The structural and geochemical data obtainedfrom the late Archean (2750–2667 Ma) Schreiber–Hemlo greenstone belt of the Superior Province,Canada, suggest that the subduction–accretion com-plex formed above a NNW-dipping subduction zone(Fig. 6). The formation of melange in this belt (Po-lat and Kerrich, in review) can be explained by arc(orogen)-parallel strike-slip faulting deforming theArchean accretionary wedge. The tectonic evolutionof the Schreiber belt is comparable to Phanero-zoic subduction–accretion complexes that evolved atoblique convergent plate boundaries (see McCaffrey,1992; Hansen, 1992; Curtis, 1997; Fuh et al., 1997).The structural characteristics of the Schreiber–Hemlo greenstone belt are similar in many respectsto those of the Vermilion district of the Wawa sub-

province to the west, and in the Abitibi greenstonebelt to the east, suggesting that oblique plate conver-gence was a continental-scale phenomenon duringthe growth of the southern Superior Province in thelate Archean (see Hudleston et al., 1988; Bauer andBidwell, 1992; Robin and Cruden, 1994; Mueller etal., 1996; Kusky and Hudleston, in press).

Accretion of immature oceanic arcs appears tohave been a major mechanism of crustal growth inArchean orogens. Reymer and Schubert (1986) andStein and Goldstein (1996), however, argued thatoceanic arc-accretion alone is insufficient to accountfor the rapid crustal growth in Precambrian shields.Furthermore, most oceanic arcs are characterizedby mafic composition (Arculus, 1981; Anderson,1982), whereas the continental crust is andesitic incomposition (Taylor and McLennan, 1995).

2.2. Ophiolite accretion

Ophiolites are a distinctive association of alloch-thonous rocks interpreted to form in a variety ofplate tectonic settings, including oceanic spreadingcenters, back-arc basins, forearcs, arcs, and otherextensional magmatic settings including those in as-sociation with plumes (e.g., Moores, 1982; Sylvesteret al., 1997). A complete ophiolite grades downwardsfrom pelagic sediments into a mafic volcanic com-plex that is generally made of mostly pillow basalts,underlain by a sheeted dike complex. These are un-derlain by gabbros exhibiting cumulus textures, thentectonized peridotite, resting above a thrust fault thatmarks the contact with underlying rock sequences.The term ‘ophiolite’ refers to this distinctive rockassociation, and should not be used in a purelygenetic way (cf. Bickle et al., 1994) to refer toallochthonous oceanic lithosphere rocks formed atmid-ocean ridges.

Very few complete Phanerozoic-like ophiolite se-quences have been recognized in Archean greenstonebelts, leading some workers to the conclusion that noArchean ophiolites or oceanic crustal fragments arepreserved (e.g., Bickle et al., 1994). However, as em-phasized by Sylvester et al. (1997), the original defi-nition of ophiolites (Anonymous, 1972) includes ‘dis-membered’, ‘partial’, and ‘metamorphosed’ varieties,and there is no justification for new arbitrary defi-nitions that attempt to exclude portions of Archean


Fig. 5. (a, b) Transposed beds, folds, shear planes, and veins within turbidites. (c) A rhyolitic melange block within mylonitized volcanicmatrix (Steel Lake area). (d) A gabbroic block in a sedimentary (pelitic) matrix (Ripple Lake area, view toward NNE). (e) An epidotized,silicified basaltic block within a sheared volcanic matrix (Middleton area, view toward SSW; top of photo to the right). (f) Sigmoidalquartz veins in volcanic rocks of the Middleton area, indicating right-lateral shearing during D2 (plan view).

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Fig. 6. Interpreted geodynamic evolution of the Schreiber–Hemlo greenstone belts, and the formation of melange. Primitive mantle-normalized trace element diagramsdiscriminating tectonically imbricated and fragmented ocean plateau tholeiites, tholeiitic to calc-alkaline arc volcanics, trench turbidites, and arc gabbros and TTG are fromPolat et al. (1998).


Fig. 7. Structural sequence columns through several Archean greenstone belt sections interpreted as ophiolites.

greenstone belts that contain two or more parts ofthe full ophiolite sequence, especially in structurallycomplex settings such as found in greenstone beltslike the Schreiber–Hemlo belt (e.g., Harper, 1985; deWit et al., 1987; Kusky, 1990, 1991). Similarly, manyProterozoic ophiolites are dismembered, or partial se-quences (Kroner, 1985; Berhe, 1990; Dann, 1991).

Archean oceanic crust was possibly thicker thanProterozoic and Phanerozoic counterparts, result-ing in accretion predominantly of the upper sec-tion (basaltic) of oceanic crust (Burke et al., 1976;Moores, 1986; Hoffman and Ranalli, 1988; Burke,1995). The crustal thickness of Archean oceaniccrust may in fact have resembled modern oceanicplateaux (e.g., Sleep and Windley, 1982). If thiswere the case, complete Phanerozoic-like MORB-type ophiolite sequences would have been very un-likely to be accreted or obducted during Archeanorogenies. In contrast, only the upper, pillow lava-dominated sections would likely be accreted.

Portions of several Archean greenstone belts havebeen interpreted to contain dismembered or partialophiolites, that are more complete than the basalt=chert slivers preserved in the Schreiber–Hemlo belt(Fig. 7). Accretion of MORB-type ophiolites hasbeen proposed as a mechanism of continental growthin a number of Archean, Proterozoic, and Phanero-

zoic orogens (Table 2). It is worthwhile to investigatethese claims to better understand the crustal structureand tectonic setting in which these Archean ophio-lites formed. Several suspected Archean ophioliteshave been particularly well-documented. One of themost disputed is the circa 3.5-Ga Jamestown ophio-lite in the Barberton greenstone belt (Fig. 7) of theKaapvaal craton (Fig. 1; Brandl and de Wit, 1997).de Wit et al. (1987) describe a 3-km-thick tectono-magmatic sequence including a basal peridotite tec-tonite unit with chemical and textural affinities toAlpine-type peridotites, overlain by an intrusive–extrusive igneous sequence, and capped by a chert–shale sequence. This partial ophiolite is pervasivelyhydrothermally altered and shows chemical evidencefor interaction with sea water with high heat andfluid fluxes (de Wit et al., 1990). SiO2 and MgOmetasomatism and black-smoker-like mineralizationis common, with some hydrothermal vents traceableinto banded iron formations, and subaerial mudpoolstructures. These features led de Wit et al. (1982,1992) to suggest that this ophiolite formed in ashallow sea, and was locally subaerial, analogous tothe plume-impinged Reykjanes ridge of Iceland. Inthis sense, Archean oceanic lithosphere may havelooked very much like younger oceanic plateauxlithosphere.

T.M. Kusky, A. Polat / Tectonophysics 305 (1999) 43–73 55Ta

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95)

Bel

ingw

egr

eens

tone

belt

(Kus

kyan

dK

idd,

1992

)K

aapv

aalc

rato

n(d

eR

onde

and

deW

it,19

94;d

eW

itet

al.,

1992

)Sl

ave

Prov

ince

(Hel

mst

aedt

etal

.,19

86;K

usky

,199

1)A

ldan

–Sta

novi

kSh

ield

(Dob

rets

ovet

al.,

1997

)Y

ilgar

ncr

aton

(Sen

gor

and

Nat

al’i

n,19

96)

Yilg

arn

crat

on(M

yers

,199

3)?

Yilg

arn

crat

on(M

yers

,199

3)Su

peri

orPr

ovin

ce(K

imur

aet

al.,

1993

)K

aapv

aalc

rato

n(B

rand

land

deW

it,19

97)

Kos

tom

uksh

agr

eens

tone

belt

(Puc

htel

etal

.,19

97a)

Pilb

ara

crat

on(M

yers

,199

3;K

iyok

awa

and

Tair

a,19

95)

Kaa

pvaa

lcra

ton

(de

Wit

etal

.,19

87;

deR

onde

and

deW

it,19

94)

Yilg

arn

crat

on(M

yers

and

Swag

er,

1997

)(A

bbot

t,19

96)

Ald

an–S

tano

vik

Shie

ld(D

obre

tsov

etal

.,19

97)

Yilg

arn

crat

on(F

ripp

and

Jone

s,19

97)

Pilb

ara

crat

on(B

arle

y,19

97)

Bur

keet

al.(

1976

)Pi

lbar

acr

aton

(Oht

aet

al.,

1996

)Ta

rney

etal

.(19

76)

Prot

eroz

oic

Bir

imia

nSh

ield

(Abo

ucha

mie

tal.,

1990

;Boh

eret

al.,

1992

)T

rans

-Hud

son

orog

en(S

tern

etal

.,19

95a)

Ara

bian

–Nub

ian

Shie

ld(K

rone

r,19

85;B

erhe

,199

0)Y

avap

ai–M

azat

zalo

roge

ns(C

ondi

e,19

86)

Ara

bian

–Nub

ian

Shie

ld(S

engo

ran

dN

atal

’in,

1996

)A

rabi

an–N

ubia

nSh

ield

(Kro

ner,

1985

;Ste

inan

dG

olds

tein

,199

6)Y

avap

ai-M

azat

zalo

roge

ns(C

ondi

e,19

86)

Bal

ticSh

ield

(Par

k,19

91)

Bal

ticSh

ield

(Par

k,19

91)?

Tra

ns-H

udso

nor

ogen

(Ste

rnet

al.,

1995

a)B

altic

Shie

ld(P

ark,

1991

)Y

avap

aior

ogen

(Dan

n,19

91)

Tra

ns-H

udso

nor

ogen

(Ste

rnet

al.,

1995

a,b)

Pan-

Afr

ican

orog

eny,

cent

ralB

razi

l(P

imen

tela

ndFu

ck,1

992)

Gre

nvill

eor

ogen

(Gar

riso

n,19

81)

Ara

bian

–Nub

ian

Shie

ld(B

erhe

,199

7)

Bir

imia

mte

rran

e(B

oher

etal

.,19

92)

Tra

ns-H

udso

n(S

t-O

nge

etal

.,19

89)

Ara

bian

–Nub

ian

Shie

ld(W

indl

eyet

al.,

1996

)Pa

n-A

fric

anor

ogen

(Bod

inie

ret

al.,

1989

)

Phan

eroz

oic

Nor

th-A

mer

ican

Cor

dille

ra(B

en-A

vrah

amet

al.,

1981

;Sam

son

and

Pach

ett,

1991

;Ric

hard

set

al.,

1991

;Con

die,

1997

a)

Sout

heas

tPac

ific

(Ham

ilton

,198

8Te

thya

nop

hiol

ites

(Moo

res,

1982

;D

ewey

,197

7;Se

ngor

,199

0;Po

late

tal

.,19

96a)

Kar

akay

aba

sin,

Tur

key

(Sen

gor

and

Yilm

az,1

981)

Ura

lides

(Ham

ilton

,197

0)

Nor

ther

nA

ndes

and

cent

ralA

mer

ica

(Sto

rey

etal

.,19

91;A

lvar

ado

etal

.,19

97)

Nor

th-A

mer

ican

Cor

dille

ra(S

amso

nan

dPa

chet

t,19

91;S

alee

by,1

992;

Mill

eret

al.,

1992

)

App

alac

hian

ophi

olite

s(D

ewey

and

Bir

d,19

71;M

elan

con

etal

.,19

97)

Roc

asV

erde

sba

sin,

Chi

le(H

anso

nan

dW

ilson

,199

1)Ja

pan

(Mat

suda

and

Uye

da,1

971)

Tari

mba

sin

(Sen

gor

etal

.,19

96)

And

esan

dC

entr

alA

mer

ica

(Tar

dyet

al.,

1994

;Ver

gara

etal

.,19

95)

Nor

th-A

mer

ican

Cor

dille

ran

ophi

olite

s(M

oore

s,19

82;C

oney

,19

89;S

alee

by,1

992)

Gal

ice

basi

n,U

.S.C

ordi

llera

(Bur

chfie

leta

l.,19

92)

Alta

ids

and

Nip

pino

ids

(Sen

gor

and

Nat

al’i

n,19

96)

Japa

n(K

imur

aet

al.,

1994

)G

amila

roit

erra

ne,A

ustr

alia

(Aitc

hiso

nan

dFl

ood,

1994

)So

uthe

astP

acifi

can

dA

ustr

alia

nop

hiol

ites

(Ham

ilton

,198

8;A

itchi

son

etal

.,19

94)

App

alac

hian

s(v

anSt

aala

ndC

olm

an-S

add,

1997

)

Qin

ling

belt,

Chi

na(X

ueet

al.,

1996

)A

ppal

achi

ans

(Den

nis

and

Wri

ght,

1997

;O’B

rien

etal

.,19

97)


Several partial or dismembered ophiolites havebeen described from the Slave Province (Figs. 1,7 and 8), although these too have been disputed(e.g., King and Helmstaedt, 1997). From the PointLake greenstone belt in the central Slave Province,Kusky (1991) described a fault-bounded sequencegrading downwards from shales and chemical sed-iments (umbers) into several kilometers of pillowlavas intruded by dikes and sills, locally into multipledike=sill complexes, then into isotropic and cumu-late-textured layered gabbro (Fig. 9). The base of thispartial Archean ophiolite sensu stricto is marked bya 1-km-thick shear zone composed predominantly ofmafic and ultramafic mylonites, with less-deformeddomains including dunite, websterite, wherlite, ser-pentinite, and anorthosite. By using down-plungeprojections and section-balancing techniques, Kusky(1991) estimated that the shear zone at the base ofthis ophiolite accommodated a minimum of 69 km ofslip. Although this still allows the ophiolite to haveformed at or near extended older continental crustthat forms the Anton terrane to the west (Kusky,1989), the actual amount of transport was proba-bly much greater. Syn-orogenic conglomerates andsandstones were deposited in several small foredeepbasins, and are interbedded with mugearitic lavas(and associated dikes), all deposited=intruded in aforeland basin setting (Kusky, 1991).

Kusky (1986, 1990) suggested that portions of theCameron and Beaulieu River greenstone belts of thesouthern Slave Province contain ophiolitic compo-nents (Figs. 7, 8 and 10). The belts are cut by nu-merous layer-parallel shear zones, but some sectionsare composed mostly of tholeiitic pillow basalts, oth-ers contain approximately equal quantities of pillowsand dikes, and a few sections consist of nearly 100%mafic dikes (Fig. 10). The bases of these green-stone belts are marked by shear zones up to 500m thick (locally containing melanges), with tectonicblocks of gabbro, mafic volcanics, peridotite, andslivers of the underlying quartzofeldspathic gneiss(with extensive mafic dike complexes, Fig. 11) andits autochthonous cover. Original relationships be-tween dikes in the basement complex and dikesin the basal parts of the greenstone belts have notbeen established, but older-generation mafic dikes donot cut intervening sedimentary sequences, nor theshear zone that separates the greenstone belt from

the basement (Fig. 10). Helmstaedt et al. (1986)describe a pillow lava sequence that grades downinto sheeted dikes and gabbro from the Yellowknifegreenstone belt (Figs. 7 and 8), but interpreted thebasal contact of the belt as an unconformity on abanded iron formation, an interpretation questionedby Kusky (1987). The dikes and pillow lavas aregeochemically similar to MORB (MacLaughlin andHelmstaedt, 1995), although Isachsen et al. (1991),and Isachsen and Bowring (1997) have shown thatthe Yellowknife greenstone belt contains several dif-ferent, and probably unrelated volcanic and sedimen-tary sequences, separated by as much as 50 Ma andspanning an age interval of 200 Ma.

Harper (1985) and Wilks and Harper (1997) de-scribe rocks of the South Pass area in the Wind RiverRange, Wyoming, as containing a dismembered meta-morphosed Archean ophiolite. This ophiolite containsall of the units of a complete ophiolite except the basalperidotite tectonite, and contacts between all units areshear zones. Cumulate textures in ultramafic rocksand gabbros are present, as are small exposures ofa sheeted dike complex. Pillow lavas are associatedwith metapelites and banded iron formation.

It has been argued (Bickle et al., 1994) thatthe paucity of well-developed sheeted dike com-plexes known from Archean greenstone belts indi-cates that they are not ophiolites. But sheeted dikesare not well-preserved in many Phanerozoic ophio-lites, especially when they are metamorphosed anddeformed to the extent that most Archean greenstonebelts are. Abbott (1996) argues that sheeted dikesare not necessarily formed in every ocean floor se-quence. Despite this, sheeted dike complexes havebeen discovered in several of the ophiolitic green-stone belts described above. Well-developed sheeteddike complexes have also been mapped in severallocations in the Kalgoorlie terrane of the Yilgarn cra-ton (Fripp and Jones, 1997). Multiple cooling unitsof dolerite, high-Mg mafic rocks, and serpentiniteare truncated at an angle between 35º and 80º by anunconformably overlying mafic volcanic breccia, pil-low breccia, and lenses of pillow lava that strike par-allel to bedding in overlying sedimentary rocks in theKanowna Lake area. Fripp and Jones (1997) inter-pret this unit as a sheeted dike complex overlain bya volcanic carapace. At the Cowan Lake Six Islandslocality (Fig. 12), Fripp and Jones (1997) describe


Fig. 8. Tectonic map of the Slave Province showing locations of the Cameron River, Point Lake and Yellowknife greenstone belts.(modified after Kusky, 1989). The ‘Yellowknife Domain’ is the area between the Yellowknife and Cameron River greenstone belts.Cross-hachured areas contain circa 2.7–4.0-Ga gneissic rocks and remnants (typically as enclaves in circa 2.5-Ga granites) of the Antonterrane; unpatterned areas represent metasedimentary rocks (dominantly graywacke turbidite sequences), intruded by younger granitoids,and the checkered pattern represents circa 2.7–2.6-Ga magmatic rocks of the Hackett River arc.

58T.M

.Kusky,

A.Polat/Tectonophysics

305(1999)

43–73

Fig. 9. Map of part of the Point Lake greenstone belt, showing ophiolitic stratigraphy (from Kusky, 1991). See Fig. 8 for location.

T.M.K

usky,A


(1999)43–73

59

Fig. 10. Map of part of the Cameron River greenstone belt, showing dike=sill complex (from Kusky, 1990). See Fig. 8 for location.

60T.M

.Kusky,

A.Polat/Tectonophysics

305(1999)

43–73

Fig. 11. Map of dike complex in granitoid basement complex of the Cameron River greenstone belt in the Upper Ross Lake area (after Kusky, 1990). See Fig. 8 for location.


Fig. 12. Map of cliff section of sheeted dike complex in the Lake Cowan section of the Kalgoorlie greenstone belt. Inset shows map ofthe Yilgarn craton, showing location of the Kalgoorlie greenstone terrane (modified after Fripp and Jones, 1997).

lherzolite and dunite that grade up into websteriteand gabbro with pyroxenite layers. These rocks areoverlain by high-Mg mafic and picritic basalts thatoccur in multiple tabular cooling units, interpretedas sheeted dikes, that exhibit both one-way andtwo-way chill margins. These are overlain by chert,silicified mudstone, shale and graywacke turbidites,which locally occur as partially assimilated xeno-liths (containing zircons) within the intrusive rocks.Fripp and Jones (1997) interpret the Lake Cowangreenstone locality to include the peridotitic lowerplutonic sequence that marks the transition zonebetween mantle and crust in ophiolite suites. Thistransition zone sequence is overlain by a sheeteddike complex, but the extrusive magmatic carapaceis omitted by faulting at this locality. Fripp and Jones(1997) note the many similarities between the Kal-goorlie ophiolites and Phanerozoic ophiolites such asthe Samail, Troodos, and Bay of Islands massifs.

Kimura et al. (1993) interpret parts of the LarderLake and Beardmore–Geraldton greenstone belts inthe Abitibi and Wabigoon subprovinces of the Supe-rior Province to include ophiolitic fragments accretedin arc environments, in a manner analogous to thesetting of the basalt=chert slivers of the Schreiber–Hemlo belt described above. The Larder Lake beltoccurs in the southern part of the Abitibi greenstonebelt, and consists of pillow basalts and banded ironformation (BIF) tectonically stacked with terrigenousturbidites. The pillow basalts and BIF are interpretedto be the upper part of an oceanic plate stratigraphy,offscraped and interdigitated with trench turbiditesin an accretionary wedge setting similar to Alaska(Kusky et al., 1997a) or Japan (Isozaki et al., 1990).Williams et al. (1991) and Kimura et al. (1993)also suggest a similar exotic origin for basalts andiron formation tectonically interleaved with terrige-nous turbidites in the Beardmore–Geraldton area in


the southern part of the Wabigoon subprovince. Inboth of these examples, the accreted trench turbiditesand ophiolitic slivers are intruded and overprintedby arc-related plutons and lavas, formed when thetrench stepped back and intruded its own accre-tionary wedge. A similar accretionary wedge settingand oceanic crustal origin for slivers of basalt ingreenstone belts of the Pilbara craton has been pro-posed by Isozaki et al. (1992).

Evidence for the creation and obduction ofoceanic crust in the Archean is not limited to field re-lationships as described above. Jacob et al. (1994) re-port that the geochemistry of diamondiferous eclog-ites from the Udachnaya Mine, Siberia (Puchtel etal., 1997b), are most consistent with derivation fromsubducted slabs of Archean oceanic crust that wereextensively hydrothermally altered prior to subduc-tion. Similarly, many eclogite samples from SouthAfrican kimberlites are also interpreted as remnantsof subducted Archean oceanic crust (e.g., Jagoutz etal., 1984; MacGregor and Manton, 1986).

In summary, dismembered ophiolites are a wide-spread component of Archean greenstone belts, andmany of these apparently formed as the upper parts ofArchean oceanic crust. Most of these appear to havebeen accreted within forearc and intra-arc tectonicsettings. The observation that Archean greenstonebelts have such an abundance of accreted ophioliticfragments compared to Phanerozoic orogens sug-gests that thick, relatively buoyant, young Archeanoceanic lithosphere may have had a rheological struc-ture favoring delamination of the uppermost parts dur-ing subduction and collisional events (Hoffman andRanalli, 1988).

2.3. Oceanic plateau accretion

Oceanic plateaux are thicker than normal oceaniccrust formed at mid-ocean ridges: they are morebuoyant and relatively unsubductable, forming po-tential sources of accreted oceanic material tothe continental crust at convergent plate bound-aries (Burke et al., 1978; Ben-Avraham et al.,1981; Richards et al., 1991; Burchfiel et al., 1992;Desrochers et al., 1993; Abbott, 1996; Abbott et al.,1997; Sengor et al., 1996). Accretion of oceanicplateaux has been proposed as a mechanism ofcrustal growth in a number of orogenic belts, includ-

ing Archean, Proterozoic, and Phanerozoic examples(Table 2). Oceanic plateaux are interpreted to formfrom plumes or plume heads that come from thelower mantle (D00) or the 670 km discontinuity, andthey may occur either within the interior of plates, orinteract with the upper mantle convective=magmaticsystem and occur along mid-ocean ridges (e.g., Kin-caid et al., 1996; review by Abbott and Mooney,1995). Storey et al. (1991) suggest that oceanicplateaux may be sites of komatiite formation pre-served in Phanerozoic through Archean mountainbelts, based on their correlation of allochthonouskomatiites and high-MgO lavas of Gorgona Island,Curacao, and in the Romeral fault zone, with theCretaceous Caribbean oceanic plateau.

Portions of several komatiite-bearing Archeangreenstone belts have been interpreted as pieces ofdismembered Archean oceanic plateaux (Kusky andKidd, 1992; Abbott and Mooney, 1995; Puchtel etal., 1997a). For instance, parts of several greenstonebelts in the southern Zimbabwe craton (Blenkinsopet al., 1997; Fig. 13) are allochthonous (Kusky andWinsky, 1995) and show a similar magmatic se-quence including a lower komatiitic unit overlain byseveral kilometers of tholeiitic pillow basalts. Kusky(1998) proposed that these represent a circa 2.7 Gaoceanic plateau dismembered during a collision be-tween the passive margin sequence developed onthe southern margin of the Zimbabwe craton andan exotic crustal fragment preserved south of thesuture-like Umtali line (Fig. 13).

The geochemistry of some of the basalts in theSchreiber–Hemlo belt suggests that they formed inan oceanic plateau setting that was later accretedto the forearc of a large juvenile arc system. Thistectonic collage later collided with other crustal frag-ments to form the core of the North American craton.Kimura et al. (1993) interpret geologic relationshipsin the Malartic–Val d’Or area of the Abitibi sub-province of the Superior Province to contain anexample of accreted oceanic plateau material. Ko-matiitic and tholeiitic magmatic suites are geochem-ically similar to oceanic plateau basalts from Gorg-ona (Kimura et al., 1993), and the tectonic slicesin the greenstone belt that contain these rocks, pre-serve early tectonic fabrics interpreted to result fromaccretion of the oceanic plateau. These rocks areintruded by a suite of calc-alkaline magmas, which


Fig. 13. Map of the Zimbabwe craton, showing distribution of oceanic plateau like greenstone belt fragments in black (modified afterKusky et al., 1997a).

Kimura et al. (1993) interpret as forming in lateextensional basins. They suggest a general model inwhich oceanic plateaux and normal oceanic crust are

accreted in arc environments, which causes a back-stepping of the subduction zone. As the accretionarycomplex grows it is overprinted by calc-alkaline


magmatism as the arc migrates through the formersubduction complex. Further magmatic and struc-tural events can be caused by late ridge subductionand strike-slip segmentation of the arc.

The geodynamic interpretation of Archean ko-matiites is however controversial; the eruption ofkomatiites are attributed to several other geodynamicsettings including mid-ocean ridges (de Ronde andde Wit, 1994), intra-oceanic plume setting (Jochumet al., 1991; Xie et al., 1993; Arndt et al., 1997), sub-duction zones (Parman et al., 1997), and continentalrifts (e.g., Nisbet et al., 1993a). Similarly, ocean floortholeiites with flat REE patterns are interpreted ei-ther as remnants of Archean mid-ocean ridge basalt(AMORB; Ohta et al., 1996), or Archean oceanicplateaus (Kimura et al., 1993; Xie et al., 1993).The presence of more-abundant komatiites, high-Mgtholeiites, and slab-derived, high-La=Ybn, high-AlTTG’s suggests slightly higher average mantle tem-peratures in the Archean than the post-Archean, re-sulting in thicker Archean oceanic crust up to 25 km(Sleep and Windley, 1982; McKenzie and Bickle,1988; Drummond and Defant, 1990; Langmuir etal., 1992; Martin, 1993; Grove et al., 1994; Condie,1997b). In addition, slightly higher mantle tempera-tures in the Archean mantle may have produced largenumbers of oceanic plateaus, up to 60–80 km thick,derived from mantle plumes (Storey et al., 1991;McDonough and Ireland, 1993; Nisbet et al., 1993b;Abbott, 1996; Arndt et al., 1997).

Average geochemical compositions of the conti-nental crust, however, are not consistent with oceanplateau accretion alone. For example, modern oceanplateaus such as Ontong Java, Kerguelen, and theNauru basin have near flat REE patterns on chon-drite-normalized diagrams (see Floyd, 1989; Salterset al., 1992; Mahoney et al., 1993), whereas the con-tinental crust has LREE-enriched patterns. Similarly,areally extensive Mg- to Fe-tholeiite sequences of theSuperior Province and the Arabian–Nubian Shield(Fig. 1, Berhe, 1997), which are interpreted as frag-ments of Archean and Proterozoic ocean plateaux,respectively, have nearly flat REE patterns (Kroner,1985; Xie et al., 1993; Desrochers et al., 1993; Polatet al., 1998).

Parts of many Archean, Proterozoic and Phanero-zoic greenstone belts interpreted as oceanic plateaufragments are overprinted by arc magmatism, sug-

gesting that they either formed the basement of intra-oceanic island arcs, or they have been intruded byarc magmas following their accretion (Abouchami etal., 1990; Desrochers et al., 1993; Condie, 1997a).Condie (1997b) suggested that the upper and lowercontinental crusts have grown through the accre-tion of oceanic island arcs and ocean plateaus, re-spectively. In summary, accreted oceanic plateauxmay form a significant component of the continentalcrust, although most are structurally disrupted andoverprinted by arc magmatism.

2.4. Back-arc basin accretion

The formation, closure, and preservation ofback-arc basin sequences has proven to be a pop-ular model for the evolution of some greenstonebelts (Table 2; see also van Staal and Colman-Sadd, 1997), especially since a direct comparisonbetween the Mesozoic Rocas Verdes of Patagoniaand Archean greenstone belts was made by Tarney etal. (1976). Despite the elegance of this model and itsapparent applicability to numerous greenstone belts,Stern and de Wit (1997) caution that the analogyis only directly applicable to those greenstone beltswith simple histories, emphasizing that the evolutionof the Rocas Verdes spans only 50 Ma.

Like some Archean greenstone belts, MesozoicRocas Verdes igneous rocks include mainly pillowbasalts, dikes, sills, and gabbros, interpreted as par-tial ophiolite sequences formed along mid-oceanridge type spreading centers (Dalziel et al., 1974;Stern and de Wit, 1997). Mafic dikes that are compo-sitionally similar to those in the ophiolitic sequencescut older, dominantly tonalitic–trondhjemitic base-ment rocks. These formed when the arc sequencewas rifted, forming a narrow basin with igneous se-quences initially fed by dikes that intruded throughthe older basement. As spreading continued, a mid-ocean ridge style spreading center developed, inwhich oceanic crust was generated. Platform andcoarse-grained arc-derived sediments were depositedon the east and west sides of the basin, whereas chertand fine-grained turbidites overlie pillow lavas in thecenter of the basin (de Wit and Stern, 1981). Closureof the back-arc basin formed folds and cleavage insedimentary rocks, and initiated major shear zonesin the basement and ophiolitic rocks. Some por-


tions of the ophiolitic sequence were thrust overa foreland basin sequence developed during basinclosure, whereas other portions of the igneous se-quence were shortened in situ and remain essentiallyautochthonous (de Wit and Ashwal, 1997).

Paradoxically, the dominance of buoyant sub-duction styles in the Archean should have led todominantly compressional arc systems, but manyworkers suggest back-arc basins (which form inextensional arcs) as a modern analog for Archeangreenstone belts. Several workers have proposed thatrocks of the ‘Yellowknife domain’ (Fig. 8) in thesouthern Slave Province may have been depositedin a back-arc basin (e.g., Kusky, 1986; MacLaugh-lin and Helmstaedt, 1995; Isachsen and Bowring,1997). Data supporting this interpretation includethe geochemical signature of metavolcanic and intru-sive rocks that is consistent with intrusion througholder continental crust (MacLaughlin and Helm-staedt, 1995), and zircons in felsic volcanic rockswith ages older than the main volcanic sequence,and approaching the age of the nearby gneissic rocks(Isachsen and Bowring, 1997). In a relationship rem-iniscent of the Mesozoic Rocas Verdes, the easternmargin of the proposed back-arc basin is marked bya tonalitic gneiss complex, cut by extensive maficdikes that are compositionally similar to those of theadjacent greenstone (ophiolitic) belts (e.g., Fig. 9;Kusky, 1990). Other back-arc basin settings havebeen proposed for components of the Sao Francisco(Baars, 1997), Aldan–Stanovik (Dobretsov et al.,1997), Kaapvaal (Brandl and de Wit, 1997), Yilgarn(Myers and Swager, 1997) and Pilbara (Tarney et al.,1976; Barley, 1997) cratons.

Many assignments of back-arc basin geodynamicsettings including that of the Yellowknife domain arebased in part on the geochemical affinities of lavasin ‘tectonic discriminant analysis’. Particular empha-sis has been placed on Nb–Ta–Th–La systematics.Recent studies have shown that the depletion of Nband Ta with respect to Th and La is not unique toarc and back-arc lavas (Klein and Karsten, 1995;Karsten et al., 1996). Thus, the depletion of Nb andTa in many Archean greenstone belts may not neces-sarily indicate arc or back-arc geodynamic settings.The diverse geochemical characteristics, particularlyin Th–Nb–La systematics, of Archean intra-oceanickomatiites and Mg- to Fe-tholeiites may imply het-

erogeneous mantle plumes (Polat et al., 1999), suchas the Iceland Plume (Hemond et al., 1993; Fur-man et al., 1995; Hards et al., 1995), rather than aback-arc geodynamic setting.

3. Arc–trench migration and accretionaryorogens; a paradigm for the Archean

Sengor and Natal’in (1996) discuss a special classof collisional mountain building that they term ‘Tur-kic-type orogeny’ (i.e., accretionary orogens), inwhich large, sub-continent-size accretionary com-plexes which magmatic arc axes have migratedthrough are built on one or two of the collidingcontinents before collision, and are later displacedby strike-slip faulting. These accretionary wedgesare typically built of belts of flysch, disrupted fly-sch and melange, and accreted ophiolites, plateaux,and juvenile island arcs (e.g. Kusky et al., 1997b).Sengor and Natal’in (1996) review the geology ofseveral Phanerozoic and Precambrian orogens, andconclude that Turkic- or accretionary-type orogeny isone of the principal builders of continental crust withtime (Table 2). Sengor and Natal’in’s observationsand model are particularly applicable to the recordof Archean granite–greenstone terranes, which typ-ically show important early accretionary phases fol-lowed by intrusion by arc magmatism (e.g., Jacksonand Cruden, 1995; Kusky and Vearncombe, 1997),possibly related to the migration of magmatic frontsthrough large accretionary complexes. In exampleslike the Schreiber–Hemlo belt and other parts ofthe Superior Province, many sub-parallel belts of ac-creted material are located between continental frag-ments that are separated by many hundreds of kilo-meters, and thus may represent large accretionarycomplexes that formed prior to a ‘Turkic-type’ col-lision. Late stage strike-slip faulting is important inthese Archean orogens, as in the Altaids and Nip-ponides, and may be partly responsible for the com-plexity and repetition of belts of similar characteracross these orogens.

Turkic- or accretionary-type orogens may alsoexperience late-stage extension associated with grav-itational collapse of the orogen, especially in associa-tion with late collisional events that thicken the crustin the internal parts of the orogen. In the Archean,


slightly higher mantle temperatures may have re-duced the possible height that mountains would havereached before the strength of deep-seated rocks wasexceeded, so that extensional collapse would haveoccurred at crustal thicknesses lower than those ofthe younger geological record (e.g., England andBickle, 1984; Kusky, 1993). This idea is discussedfurther in the following section.

4. Late-stage granites and cratonization

Granite–greenstone terranes are ubiquitously in-truded by late- to post-kinematic granitoid plutons,which may play a role in or be the result of someprocess that has led to the stabilization or ‘cratoniza-tion’ of these terranes (e.g., Kusky, 1993, 1998;Abbott et al., 1997). Most granite–greenstone ter-ranes also have a thick mantle root or tectosphere(Jordan, 1988; Abbott et al., 1997), characterized bya refractory composition (depleted in a basaltic com-ponent), relatively cold temperatures, high flexuralrigidity, and high shear wave velocities (Ballard andPollack, 1987; Grand, 1987, 1994; Bechtel et al.,1990).

Outward growth and accretion in granite–green-stone terranes provides a framework for the suc-cessive underplating of the lower parts of depletedslabs of oceanic lithosphere (Helmstaedt and Doig,1975; Helmstaedt and Shulze, 1989), particularly ifsome of the upper sections of oceanic crust are off-scraped and accreted, to be preserved as greenstonebelts, or eroded to form belts of graywacke tur-bidites. These underplated slabs of depleted oceaniclithosphere will be cold and compositionally buoyantcompared to surrounding asthenosphere (providingthat the basalt is offscraped and not subducted andconverted to eclogite) and may contribute to the for-mation of cratonic roots (Abbott, 1991). One of themajor differences between Archean and younger ac-cretionary orogens is that Archean subducted slabswere dominantly buoyant, whereas younger slabswere not (e.g., Abbott, 1991). This may be a re-sult of the changing igneous stratigraphy of oceaniclithosphere, resulting from a reduction in heat flowwith time, perhaps explaining why Archean cratonshave thick roots and are relatively undeformablecompared to their younger counterparts. Geometric

aspects of underplating these slabs predict that theywill trap supra-subduction mantle wedges of morefertile and hydrated mantle (Kusky, 1993, 1986),from which later generations of basalt can be gener-ated.

Many granites in Archean terranes appear to beassociated with crustal thickening and anatexis dur-ing late stages of collision. However, some late-stagegranitoids may be a direct result of decompressionalmelting associated with upper-crustal extensionalcollapse of Archean orogens thickened beyond theirlimit to support thick crustal sections, as determinedby the strength of deep-seated rocks (e.g., Dewey,1988). Decompressional melting generates basalticmelts from the trapped wedges of fertile mantle,which intrude and partially melt the lower crust. Themelts assimilate lower crust, become more silicicin composition, and migrate upward to solidify inthe mid to upper crust, as the late- to post-kine-matic granitoid suite (e.g., Nelson, 1991; Kusky,1993). In this model, the tectosphere (or mantle root)becomes less dense (compositionally buoyant) andcolder than surrounding asthenosphere, making it astable cratonic root that shields the crust from furtherdeformation.

Late-stage strike-slip faults that cut many Archeancratons may also play an important role in craton sta-bilization (e.g., Kusky, 1998). Specifically the steepshear zones may provide conduits for massive fluidremobilization and escape from the subcontinentallithospheric mantle, which would both stabilize thecratonic roots of the craton, and initiate large-scalegranite emplacement into the mid and upper crust.

5. Conclusions: continental growth ataccretionary convergent margins

Although the rate of continental growth is a matterof geological debate (Fyfe, 1978; Dewey and Wind-ley, 1981; Armstrong, 1991; McCulloch and Bennett,1994; Taylor and McLennan, 1995), most geologicaldata indicate that the continental crust has grown byaccretionary and magmatic processes taking place atconvergent plate boundaries since the early Archean(Burke et al., 1976; Sleep and Windley, 1982; Friendet al., 1988; Card, 1990; Condie, 1994, 1997a; Tay-lor and McLennan, 1995; Sengor and Natal’in, 1996;


Windley et al., 1996). Arc-like trace-element charac-teristics of continental crust suggest that subductionzone magmatism has played an important role in thegeneration of the continental crust (Hofmann, 1988;Tarney and Jones, 1994; Rudnick, 1995; Taylor andMcLennan, 1995).

Convergent margin accretionary processes thatcontribute to the growth of the continental crust canbe divided into five major groups (Table 2): (1)oceanic plateau accretion; (2) oceanic island arc ac-cretion; (3) normal ocean crust (mid-ocean ridge)accretion=ophiolite obduction; (4) back-arc basin ac-cretion; and (5) arc–trench migration=Turkic-typeorogeny accretion. We have illustrated these pro-cesses using specific well-studied examples of each,emphasizing the Schreiber–Hemlo area of the Su-perior Province because it shows features related toseveral of these process. These early accretionaryprocesses are typically followed by intrusion of late-stage anatectic granites, late gravitational collapse,and late strike-slip faulting. Together, these processesrelease volatiles from the lower crust and mantle andhelp to stabilize young accreted crust and form stablecontinents.

Isotopic data from numerous Archean and post-Archean orogenic belts are consistent with the ac-cretion of juvenile material, rapid and episodiccrustal growth (Reymer and Schubert, 1986; Sam-son and Pachett, 1991; Boher et al., 1992; Taylorand McLennan, 1995; Stein and Goldstein, 1996;and references therein). Stein and Hofmann (1994)proposed that episodic crustal growth and major oro-genic events may have been associated with majorplume activities throughout earth history. They de-fined this relationship as mantle overturn and majororogenies (MOMO). This model can explain global,episodic accretion of Archean komatiite–tholeiiticbasalt sequences. However, the abundance of syn-to post-kinematic TTG in the Archean upper con-tinental crust, rather than ultramafic–mafic volcanicsequences (Taylor and McLennan, 1995), requiresthe production of voluminous slab, mantle wedge,and lower crustal-derived calc-alkaline arc magmasas complementary to plateau accretion (Drummondand Defant, 1990; Feng and Kerrich, 1992; Martin,1993).

Sengor et al. (1993) and Sengor and Natal’in(1996) proposed a new type of orogeny, so-called

‘Turkic or accretionary-type orogeny’, for continen-tal growth. These orogenic belts possess very largesutures (up to several hundred km wide) charac-terized by subduction–accretion complexes and arc-derived granitoid intrusions, similar to the Circum-Pacific accreted terranes (e.g., Alaska, Japan). Thesesubduction–accretion complexes are composed oftectonically juxtaposed fragments of island arcs,back-arc basins, ocean islands=plateaux, trench tur-bidites, and micro-continents (Sengor, 1993; Sengorand Natal’in, 1996). Another important feature ofthese orogens is the common occurrence of orogenparallel strike-slip fault systems, resulting in lateralstacking and bifurcating lithological domains (Sen-gor and Natal’in, 1996). In these respects, the accre-tionary-type orogeny may be considered as a unifiedaccretionary model for the growth of the continentalcrust.

Acknowledgements

We thank Steve Marshak, Peter Hudleston, JohnPercival, and Celail Sengor for thoughtful reviewsof the manuscript. Kusky acknowledges support byNSF grant 9706699, and Polat acknowledges sup-port from an NSERC grant (part of NSERC-IOR toMITEC (CAMIRO) 93EOR) awarded to R. Kerrich.

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Growth of granite–greenstone terranes at convergent margins, and stabilization of Archean cratons

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