Gondwana Research 11 (2007) 69–91 www.elsevier.com/locate/gr

Field and geochemical characteristics of the Mesoarchean (∼3075Ma) Ivisaartoq greenstone belt, southern West Greenland: Evidence for seafloor hydrothermal alteration in supra-subduction oceanic crust ⁎ Ali Polat a, , Peter W.U. Appel b, Robert Frei c, Yuanming Pan d,Yıldırım Dilek e, Juan C. Ordóñez-Calderón a, Brian Fryer a,f, Julie A. Hollis b, Johann G. Raith g

a Department of Earth Sciences, University of Windsor, Windsor, ON, Canada N9B 3P4 b Geological Survey of Denmark and Greenland, 1350 Copenhagen, Denmark c Geological Institute, University of Copenhagen, 1350 Copenhagen, Denmark d Department of Geological Sciences, University of Saskatchewan, Saskatoon, SK., Canada S7N 5E2 e Department of Geology, Miami University, Oxford, OH 45056, USA f Great Lakes Institute for Environ. Res., University of Windsor, Windsor, ON, Canada N9B 3P4 g Department of Applied Geosciences and Geophysics, University of Leoben, Leoben, A-8700, Austria Received 12 December 2005; received in revised form 12 February 2006; accepted 14 February 2006

Abstract

The Mesoarchean (ca. 3075Ma) Ivisaartoq greenstone belt in southern West Greenland includes variably deformed and metamorphosed pillow basalts, ultramafic flows (picrites), serpentinized ultramafic rocks, gabbros, sulphide-rich siliceous layers, and minor siliciclastic sedimentary rocks. Primary magmatic features such as concentric cooling-cracks and drainage cavities in pillows, volcanic breccia, ocelli interpreted as liquid immiscibility textures in pillows and gabbros, magmatic layering in gabbros, and clinopyroxene cumulates in ultramafic flows are well preserved in low-strain domains. The belt underwent at least two stages of calc-silicate metasomatic alteration and polyphase deformation between 2963 and 3075Ma. The stage I metasomatic assemblage is composed predominantly of epidote (now mostly diopside)+quartz+plagioclase±hornblende ±scapolite, and occurs mainly in pillow cores, pillow interstitials, and along pillow basalt-gabbro contacts. The origin of this metasomatic assemblage is attributed to seafloor hydrothermal alteration. On the basis of the common presence of epidote inclusions in diopside and the local occurrence of epidote-rich aggregates, the stage I metasomatic assemblage is interpreted as relict epidosite. The stage II metasomatic assemblage occurs as concordant discontinuous layered calc-silicate bodies to discordant calc-silicate veins commonly associated with shear zones. The stage II metasomatic assemblage consists mainly of diopside+garnet+amphibole+plagioclase+quartz±vesuvianite±scapolite±epidote±titanite ±calcite±scheelite. Given that the second stage of metasomatism is closely associated with shear zones and replaced rocks with an early metamorphic fabric, its origin is attributed to regional dynamothermal metamorphism. The least altered pillow basalts, picrites, gabbros, and diorites are characterized by LREE-enriched, near-flat HREE, and HFSE (especially Nb)-depleted trace element patterns, indicating a subduction zone geochemical signature. Ultramafic pillows and cumulates display large positive initial εNd values of +1.3 to +5.0, consistent with a strongly depleted mantle source. Given the geological similarities between the Ivisaartoq greenstone belt and Phanerozoic forearc ophiolites, we suggest that the Ivisaartoq greenstone belt represents Mesoarchean supra-subduction zone oceanic crust. © 2006 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

Keywords: Greenstone belt; Archean; Pillow basalt; Picrite; Epidosite; Ocelli; Forearc; Ophiolite

1. Introduction

Archean greenstone belts are the product of multiple ⁎ Corresponding author. geological processes such as tectonism, magmatism, sedimen- E-mail address: [email protected] (A. Polat). tation, and metamorphism, operating over different spatial and

1342-937X/$ - see front matter © 2006 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2006.02.004 70 A. Polat et al. / Gondwana Research 11 (2007) 69–91 temporal scales (Condie, 1981; de Wit, 1998; Polat et al., 1998; Hall et al., 1987; Chadwick, 1990; Appel, 1997). The presence van Kranendonk, 2004; Condie, 2005; Polat and Kerrich, 2006; of these primary structures provides a unique opportunity to Benn et al., 2006). Geochemical data accumulated from studies study the characteristics of hydrothermal alteration, physical of Archean greenstone belts over the last three decades are volcanology, and geodynamic processes. This study has been consistent with diverse types of volcanic rocks on all major made possible by the detailed mapping and recognition of Archean cratons, reflecting diverse geodynamic processes pillow basalts in the Ivisaartoq greenstone belt in the 1970s and (Dostal and Mueller, 1997; Polat et al., 1998; Kusky and 1980s (Friend and Hall, 1977; Friend et al., 1981; Brewer et al., Polat, 1999; Polat and Hofmann, 2003; Dostal and Mueller, 1984; Chadwick, 1985, 1986; Hall et al., 1987; Chadwick, 2004; Manikyamba et al., 2005; Smithies et al., 2005a,b; and 1990). The objectives of this study are two-fold: (1) to assess the references therein). These studies have documented two major geochemical and mineralogical effects of the post-magmatic types of volcanic rock associations: (1) an oceanic plateau hydrothermal alteration on pillow basalts, gabbros, diorites, and association composed of compositionally relatively uniform ultramafic pillows and cumulates; and (2) to understand the komatiites and Mg- to Fe-rich tholeiitic basalts erupted from geodynamic setting of the Ivisaartoq greenstone belt in which mantle plumes and (2) a compositionally diverse intra-oceanic the mafic-ultramafic rocks experienced hydrothermal alteration. island arc association, dominated by ‘normal’ tholeiitic to calc- alkaline basalts, andesites, dacites, and rhyolites (BADR). In 2. Regional geology and field characteristics addition, the latter association includes small volumes of boninites, picrites, low-Ti tholeiites (LOTI), adakites, high- The Ivisaartoq greenstone belt contains the largest magnesian andesites (HMA), and Nb-enriched basalts (NEB) Mesoarchean supracrustal assemblage in southern West Green- (Polat and Kerrich, 2004, 2006). land (Fig. 1; Hall and Friend, 1979; Brewer et al., 1984; Despite the two major phases of deformation, including Chadwick, 1985, 1986, 1990; Friend and Nutman, 2005a). It is isoclinal folding, and amphibolite-facies metamorphism, pillow located in the central part of the inner Godthåbsfjord of the structures, volcanic breccia, cumulate and liquid immiscibility northeastern Nuuk region (Fig. 1). The belt occurs within the textures, as well as magmatic layering have been well preserved recently recognized Mesoarchean (∼3075–2950Ma) Kapisilik in low-strain domains of the Mesoarchean Ivisaartoq greenstone tectonic terrane (Friend and Nutman, 2005a). The Kapisilik belt (Friend et al., 1981; Hall, 1981; Chadwick, 1985, 1986; terrane is tectonically bounded by the Paleoarchean Isukasia

Fig. 1. A simplified geological map of the northeastern Nuuk region, showing the tectonic terranes of late to early Archean age and location of the Ivisaartoq belt. Modified after Friend and Nutman (2005a). A. Polat et al. / Gondwana Research 11 (2007) 69–91 71 terrane (3600–3800Ma) to the north, and by the Paleoarchean Færingehavn and the Neoarchean Tre Brødre terranes to the south to west, respectively (Fig. 1; Friend and Nutman, 2005a). The Isukasia and Færingehavn terranes may represent the fragments of the same Paleoarchean continent that rifted apart at about 3500Ma and onward, resulting in the opening of an oceanic basin. The subsequent closure of this ocean in the late Archean may have resulted in the amalgamation of the above tectonic terranes as oceanic island arcs and/or continental blocks. The Kapisilik and Isukasia terranes were juxtaposed (collided) and metamorphosed by 2950Ma. It appears that the collision between the southern Færingehavn and the Kapisilik terranes took place at about 2800Ma (Friend and Nutman, 2005a). Field relationships indicate that the Isukasia terrane is structurally overlain by the Kapisilik terrane to the south; and the Kapisilik terrane is in turn structurally overlain by the Færingehavn and Tre Brødre terranes to the south–southwest. The precise age of the Ivisaartoq greenstone belt is unknown. Siliceous volcano-sedimentary rocks have yielded an average U–Pb zircon age of 3075Ma (Friend and Nutman, 2005a), constraining the maximum age of the belt. The Ivisaartoq supracrustal rocks are intruded by weakly deformed 2961 ±12Ma granites to the north, constraining the minimum age of the belt (Chadwick, 1990; Friend and Nutman, 2005a). The Ivisaartoq rocks are truncated by an up to 2-m-thick mylonite zone to the south, separating the belt from an association of leucogabbros and anorthosites. The leucogabbro and anortho- site association is intruded by 2963±8Ma old tonalites and granodiorites (now gneisses). On the basis of field observations and zircon ages, Friend and Nutman (2005a) interpreted the Fig. 2. A simplified tectonostratigraphic column of the Ivisaartoq belt. Modified after Chadwick (1986, 1990). mylonite zone as a post-2960Ma structure deforming the Kapisilik terrane. The Ivisaartoq greenstone belt is composed mainly of The foliation is cut by 1- to 10-cm-thick deformed (folded) quartz metamorphosed mafic to ultramafic volcanic rocks, gabbros, veins. Up to 50m wide and 5km long rusty layers of pyrite-rich minor diorites, and ultramafic rocks (Fig. 2; Hall, 1981; felsic schist are exposed discontinuously inthe lower unit (Fig.2). Chadwick, 1985, 1986, 1990). Metasedimentary rocks consti- Discordanttoconcordant,relativetodominantfoliation,veinsand tute a volumetrically minor component of the belt (Fig. 2). calc-silicate bodies containing diopside+garnet+hornblende Volcanic rocks consist dominantly of deformed pillow basalts occur mainly near the top of the magnetic marker and in the and ultramafic lava flows (Chadwick, 1990; this study). The lowermost part of the upper amphibolite unit (Fig. 3a–c). The Ivisaartoq supracrustal sequence is subdivided into a lower and thickness of these calc-silicate rocks ranges from several an upper amphibolite unit (Fig. 2). These units are separated by centimeters to several meters (Fig. 3a–c). a thin layer (up to 50m thick) of magnetite-rich ultramafic The upper group is composed mainly of variably deformed schists, called the “magnetic marker” (Fig. 2; Chadwick, 1986, pillow basalts, volcanic breccia, foliated amphibolite, actinolite 1990). Hydrothermal alteration of the magnetic marker and schist, gabbro, diorite, ultramafic cumulate, serpentinite (ultra- volcanic rocks in its vicinity resulted in the formation of calc- mafic layers), and metasedimentary rocks (Figs. 2–5). Strongly silicate rocks hosting strata-bound scheelite mineralization serpentinized ultramafic rocks are exposed discontinuously as (Appel, 1994, 1997). Similar rock types in the 2.7Ga Belingwe three major layers throughout the sequence (Fig. 2; Chadwick, greenstone belt, Zimbabwe, have been interpreted as mineral- 1986, 1990). They form boudinaged ultramafic bodies up to ized high-strain zones (Hofmann and Kusky, 2004). The 600m long and up to 200m thick separated by pegmatites or intensity of deformation appears to increase towards the pillow basalts (Chadwick, 1986). Chadwick (1986) reports the boundary between the two amphibolite units, suggesting that presence of olivine in these ultramafic bodies. According to they are tectonically juxtaposed. Chadwick (1986, 1990), the protoliths of the serpentinized Volcanic rocks in the lower group are more intensely deformed peridotites intruded as sills into submarine lavas. They grade than those in their upper counterpart. They display a well- into layered gabbro with a spotty texture of hornblende grains developed foliation characterized by amphibole-rich domains. set in a plagioclase matrix (Fig. 5b). Pillow basalt structures are rare. Volcanic breccias are composed Pillow basalts are characterized by well-preserved core and ofpillowfragmentslocallywithpossibleocelliandhyaloclastites. rim structures (Figs. 3d–f and 4a–c). The least deformed pillow 72 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Fig. 3. Field photographs of the Ivisaartoq calc-silicate silicate rocks and pillow basalts. (a) A diopside+garnet+hornblende+quartz±epidote vein. (b) Actinolite schist (dark) replaced by a massive layered calc-silicate rock. (c) Boudinaged calc-silicate layers in banded amphibolite. (d) Pillow basalts with well-preserved core and rim structures. (e) Pillow basalts with concentric cooling cracks filled with quartz. (f) Pillow basalt cores and interstitials filled with quartz±pyrite. basalts have concentric cooling cracks filled mainly with quartz, Pillow rims often display silica alteration. In some locations and display way-up directions (Fig. 3e). Pillow cores are pillows have been completely silicified (Fig. 4d). A similar calc- mineralogically zoned (Fig. 3d). Many inner pillow cores silicate alteration occurs in pillow interstitials, replacing the display drainage cavities at the centre, which are either empty or interstitial hyaloclastite (Fig. 3f). No ocellus has been observed filled with quartz (Figs. 3f and 4a). The pillow cores often in the pillow rims. Some pillows are composed predominantly consist chiefly of white ellipsoidal (eye-shaped) mm- to cm- of actinolite, consistent with an ultramafic composition. sized ocelli set in a dark green finer-grained mafic matrix (Fig. Primary magmatic textures, such as cumulates, are locally 4b,c). The long axes of the ocelli are parallel to those of preserved in ultramafic flows of low-strain domains (Fig. 5a). stretched pillows, suggesting that they formed as spherical There are minor, up to several meter-thick lenses of siliciclastic structures and were later flattened by deformation (Fig. 4b,c). metasedimentary rocks in the upper amphibolites (Fig. 4f). The contacts between ocelli and matrix are sharp (Fig. 4b). In Contacts between volcanic and sedimentary rocks are sharp many cores the ocelli-matrix texture has been partly to (Fig. 4f). completely replaced by a calc-silicate metasomatic assemblage Gabbros and minor diorites occur as one to several tens of (Fig. 3d). It appears that the metasomatic assemblage grew meter-thick sills and dykes in pillow basalts (Fig. 5b). In some outward from the pillow drainage cavities towards the rims. places sills occur between pillow basalts and ultramafic flows. A. Polat et al. / Gondwana Research 11 (2007) 69–91 73

Fig. 4. (a) A pillow core drainage cavity partly filled with quartz. (b) Ocelli within a pillow cores. (c) Ocelli-rich pillow core and quartz-filled cavities in the pillow interstitials. (d) Silicified pillow rims. (e) Deformed pillow breccia and hyaloclastite. (f) Metasediment (felsic gneiss) and amphibolite contact.

Chilled margins between pillow basalts and gabbroic dykes are quartz (20–30%)+plagioclase (20–30%)+epidote (0–10%)+ preserved in several locations. Primary igneous textures and amphibole (0–10%). This assemblage occurs around empty minerals are locally preserved in low-strain domains (Fig. 5b). drainage cavities and cavities filled with quartz (Figs. 3d,f and Some gabbros also contain deformed ocelli up to 15-cm long 4c). Epidote in the pillow cores occurs most commonly as (Fig. 5c,d). The long axes of these ocelli are parallel to the inclusions (with titanite and pyrite) in the cores of diopside stretching lineations in the neighbouring amphibolites (Fig. 5c). grains (Fig. 6a). Locally, epidote-rich aggregates with abundant Like pillow basalts, gabbros and diorites underwent calc-silicate plagioclase and quartz have also been found as isolated patches metasomatic alteration mainly along fractures and pillow basalt or pods in the pillow cores (Fig. 6b). Granoblastic grains of contacts (Fig. 5e,f). epidotes in those aggregates show complex zonation (Fig. 6b). Diopside is locally present as a minor phase intergrowth with 3. Petrography epidotes in those aggregates. Other minor to trace minerals present in those epidote-rich aggregates include titanite and For descriptive purposes, the pillow cores are divided into pyrite (Fig. 6b). inner and outer parts (Fig. 3d). The inner pillow cores Foliated amphibolites are composed mainly of hornblende+ commonly contain the stage I calc-silicate metasomatic plagioclase+quartz±diopside±epidote±titanite±sulphide. assemblage composed dominantly of diopside (30–40%)+ The stage II calc-silicate formation, occurring as discordant to 74 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Fig. 5. (a) Clinopyroxene cumulates in an ultramafic flow (or shallow sill). (b) Magmatic layering in gabbros. (c–d) Flattened cm-sized ocelli in gabbros. (e) Calc- silicate alteration at a pillow basalt gabbro contact and in pillow cores. (f) Metasomatized gabbro. concordant (relative to the dominant foliation) veins and bodies grained hornblende+quartz+plagioclase+epidote. In some within amphibolites and magnetic marker, is composed outcrops, pillows have been completely replaced by quartz dominantly of a diopside+garnet+amphibole+plagioclase+ and plagioclase (Fig. 4d). quartz±vesuvianite±scapolite±epidote±titanite±calcite± Ultramafic cumulates are composed primarily of altered scheelite±K-feldspar assemblage (Fig. 3a–c). Granoblastic clinopyroxene phenocrysts (Fig. 5a). There are relict clinopyr- grains of epidote commonly occur in mutual contacts with oxene and plagioclase in some gabbros. Gabbros and diorites diopside (Fig. 6c). are composed mainly of hornblende+plagioclase±epidote ± Ocelli in the outer cores (Fig. 4b, c) consist mainly of quartz (Fig. 5b). Like those in pillow basalts, the big ocelli in plagioclase (30–50%)+quartz (30–40%)+amphibole (10– gabbros are composed dominantly of plagioclase and minor 20%)±epidote (0–5%). No internal structure has been quartz (Fig. 5c,d). observed in the ocelli. The darker matrix surrounding the ocelli is made of amphibole (50–60%)+plagioclase (20–30)+ 4. Analytical procedure and data presentation quartz (10–20%)±epidote (0–5%)±titanite (0–5%). Relict grains of clinopyroxene (cpx) have been observed in only one Compositions of epidotes and associated minerals in calc- sample, indicating that amphibole resulted from the alteration silicate assemblages from veins in the amphibolites and in the of clinopyroxene. The pillow rims are composed of fine- pillow cores (Table 1) have been determined by use of a JEOL A. Polat et al. / Gondwana Research 11 (2007) 69–91 75

30–60s, and mineral standards (almandine garnet, Al and Fe; diopside, bustamite, Mn; celestite, Sr; chromite, Cr; Ca, Mg and Si; fluorite, F; jadeite, Na; rutile, Ti; sanidine, K; and tugtupite, Cl). All whole-rock samples were powdered using an agate mill in the Department of Earth Sciences, University of Windsor, Canada. Major elements were determined by Thermo Jarrel-Ash ENVIRO II ICP at ACTLABS in Ancaster, Canada. Samples were mixed with a flux of lithium metaborate and lithium tetraborate, and fused in an induction furnace. Molten melt was immediately poured into a solution of 5% nitric acid containing an internal standard, and mixed continuously until completely dissolved. Totals of major element oxides are 100±1wt.% and the analytical precisions are 1–2%. Samples were analyzed for REE, HFSE, LILE, and transition metals (Ni, Co, Cr, and V) by a high-sensitivity Thermo Elemental X7 ICP-MS in the Great Lakes Institute for Environ. Res. (GLIER), University of Windsor, Canada, using the protocols of Jenner et al. (1990). Wet chemical procedures were conducted under clean lab conditions, and all acids were distilled twice. Approximately 100–130 mg of sample powder was used for dissolution. Samples were dissolved in a concentrated HF–HNO3 mixture at a temperature of ∼120°C for 4 days, and then further attacked with concentrated HNO3 until no residue was visible. BHVO- 1 and BHVO-2 were used as international reference materials to estimate precision and accuracy (Table 2). Analytical precisions are estimated as follows: 1–10% for REE, Rb, Li, Cs, Sr, Ba, Y, Cu, Zn, and Pb; 10 to 20% for Zr, V, Cr, Ni, Co, Nb, and U; and 20–30% for Ta and Th (Table 2). The analytical procedure for Sm–Nd isotope analyses is presented in (Frei et al., 2004). Selected elements are normalized to primitive mantle (pm) (Hofmann, 1988) and chondrite (cn) (Sun and McDonough, 1989). Nb/Nb⁎,Zr/Zr⁎ and Eu/Eu⁎ ratios, representing anomalies, were calculated with respect to the neighbouring immobile elements, following the method of Taylor and McLennan (1985). Samples were recalculated to 100% anhydrous for inter-comparisons. Mg-numbers (%) were calculated as the molecular ratio of Mg/(Mg+Fe2+), where Fe2+ is assumed to be 90% of total Fe. Zircon U–Pb age dating was carried out in the Geological Fig. 6. Backscattered electron images illustrating textural relationships between epidotes and associated minerals in calc-silicate assemblages. (a) Stage I Survey of Denmark and Greenland (GEUS) in Copenhagen epidote-rich aggregate with abundant plagioclase (Pl) and minor amounts of using laser ablation–magnetic sectorfield–inductively coupled quartz (Qtz), titanite and pyrite (Py) in a pillow core (485441). (b) Stage I plasma–mass spectrometry (LA-SF-ICP-MS). The LA-SF-ICP- epidote-rich aggregate with abundant plagioclase and quartz and minor amounts MS system consists of a New-Wave Research/Merchantek of titanite and pyrite in a pillow core (485446); note sharp boundaries between UP213 laser ablation system equipped with a frequency Fe-poor clinozoisite (Cz) and epidotes (including Fe-rich clinozoisite). (c) Stage II epidote (Ep)+clinopyroxene (Cpx)+quartz (Qtz)+titanite (Tnt) in a calc- quintupled Nd-YAG laser emitting at a wavelength of 213nm silicate vein (sample IVCS3). coupled to an Element2 (ThermoFinnigan, Bremen) single- collector double focusing magnetic sectorfield ICP-MS equipped with a fast fieldregulator for increased scanning JXA-8600 superprobe equipped with three automated wave- speed. The raw data were exported in ASCII format and length dispersive spectrometers, in the Department of Geo- processed using in-house data reduction spreadsheets. Final age logical Sciences, University of Saskatchewan. Analytical calculations were also done using in-house age-calculation conditions included an accelerating voltage of 15kV, a beam spreadsheets. The computer program IsoplotEx v. 3.0 (Ludwig, current of 50nA, beam diameters of 2–5μm, counting times of 2003) was used to carry out the final computation of ages. 76 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Table 1 Representative compositions of epidotes and associated minerals in calc-silicate assemblages a Calc-silicates in pillow cores (Stage I metasomatism) Calc-silicate veins in amphibolites (Stage II metasomatism) Mineral Epidote Scapolite Epidote Clinopyroxene Garnet Amphibole Plagioclase Sample 485440 485441 485446 IVS53 IVS54 IVS53 IVS54 IVS54 IVS54 IVS54

SiO2(wt%) 37.8 39.3 40.3 40 38.4 57.1 38.1 38.6 52 52.4 37.5 38.7 48.3 55.3 45.5 51.7 TiO2 0.11 0 0 0 0 0 0.13 0 0 0 1.03 1.16 0.37 0 0 0 Al2O3 25.2 28.5 32.2 31.1 28.3 25.6 24 28 0 1.92 12.4 14.92 9.79 1.2 34.1 29.2 Cr2O3 1.27 0.01 0.03 0.02 0 0.01 0.03 0 0.04 0.06 0.08 0.32 0.02 0.14 nd nd a Fe2O3 8.87 5.48 1.39 2.51 6.3 0.17 11.8 7.14 15 8.71 0.16 0.03 0.01 0.18 FeO 12.1 8.75 14.2 12.8 MnO nd 0.09 0.01 0.03 0.05 0.02 0.03 0.04 0.98 0.45 0.12 0.03 0.32 15 nd nd MgO nd 0.03 0.02 0.02 0.03 0 0.07 0.08 10.1 12.6 0.67 0.62 10.9 0.55 nd nd CaO 0.02 23.9 24.5 24.1 24 9.92 22.9 23.7 23.8 23.6 32.7 34.93 12.1 12.5 19.1 13.5 SrO 0.13 nd nd nd nd nd nd nd nd nd nd nd nd nd 0 0 BaO 23.6 nd nd nd nd nd nd nd nd nd nd nd nd nd 0 0

Na2O 0 0 0 0 0 5.34 0 0 0 0 0.51 0.51 0.54 0 0.63 3.84 K2O 0 0 0 0 0 0.33 0 0 0 0 0 0 0.32 0.18 0 0.03 Cl nd nd nd nd nd 2.39 nd nd nd nd nd nd nd nd nd nd –O=Cl 0.54 Total 97 97.2 98.4 97.9 97 100.3 97.1 97.5 99 99.9 99.9 99.9 97 97.7 99.3 98.5 Si 3.003 3.058 3.053 3.063 3.002 3.041 3.009 2.006 1.968 2.972 3.005 7.088 7.952 2.12 2.391 Ti 0.007 0 0 0 0 0.007 0 0 0 0.061 0.068 0.041 0 0 0 Al 2.359 2.613 2.873 2.806 2.607 2.264 2.579 0 0.085 1.156 1.367 1.693 0.204 1.87 1.59 Cr 0.08 0.001 0.002 0.001 0 0.002 0 0.001 0.001 0.005 0.02 0.002 0.016 Fe3+ 0.529 0.32 0.079 0.144 0.369 0.713 0.42 0.893 0.51 0.012 0.003 0 0.006 Fe2+ 0.389 0.274 1.733 1.533 Mg 0.002 0.01 0.001 0.003 0.006 0.004 0.005 0.58 0.706 0.045 0.041 2.386 3.216 Mn 0.009 0.002 0.002 0.001 0.002 0.004 0.005 0.032 0.014 0.014 0.004 0.04 0.067 Ca 2.011 1.996 1.991 1.98 2.014 1.965 1.982 0.985 0.951 2.775 2.909 1.906 1.928 0.953 0.668 Na 0 0 0 0 0 0 0 0 0 0.079 0.076 0.153 0 0.057 0.344 K 0 0 0 0 0 0 0 0 0 0.062 0.035 0 0.002 XFe 59 34 8 15 38 73 42 − XFe in epidote-group minerals (clinozoisite and epidote)=100×Fe/(Al+Fe2). nd=not determined. a Total iron as Fe2O3 in epidotes, garnets and plagioclase; Fe2O3 in amphiboles calculated on the basis of S(Si+Ti+Al+Cr+Fe+Mn+Mg)=13.0 and 23 oxygens.

Detailed description of this analytical method is given in Epidotes in the Stage II calc-silicate rocks are characterized Janoušek et al. (2006). by irregular patterns of compositional zonation (Fig. 6c) and vary from clinozoisite to epidote (i.e., XFe from 42 to 73; Table 5. Mineral chemistry of calc-silicate assemblages 1). Associated clinopyroxene is diopside with varying amounts of the hedenbergite component (24–45mol%), whereas other Epidotes in the stage I calc-silicate assemblages from the constituents (Al2O3, MnO and Na2O) are generally minor pillow cores also vary from clinozoisite to epidote but have (Table 1). Garnet is dominated by the grossular and andradite notably lower XFe contents (8–59; Table 1) than their stage II solid solution, with only trace amounts of pyrope and counterparts. It is noteworthy that epidote inclusions in diopside spessartine components (Table 1). Several of the large garnets are compositionally indistinguishable from those in the isolated show a pronounced zoning with a greenish fairly homogenous patches or pods (Fig. 6b,c). Sharp boundaries between Fe-poor grossularite in the centre surrounded by fine scale concentric clinozoisite and Fe-rich clinozoisite-epidote (Fig. 6c) indicate a andradite. Calcic amphiboles vary from actinolite to magnesio- compositional gap at XFe between 15 and 34, similar to the hornblende (Table 1). Plagioclase is mostly anorthite (An=88– miscibility gap in epidote-group minerals reported from other 96mol%), except for a few intermediate compositions at 66– hydrothermal and low-temperature metamorphic terranes 68mol% An. (Grapes and Hoskin, 2004 and references therein). Epidotes in one pillow core sample are also characterized by a minor 6. Geochemical results amount of Cr (up to 2.3wt.% Cr2O3). Clinopyroxene in calc- silicate assemblages in the pillow cores is somewhat richer in 6.1. Pillow basalts the diopside component (65–80mol%) than its vein counter- part. Calcic amphibole is invariably actinolite (≤4.0wt.% The altered inner pillow cores have relatively low MgO Al2O3). Plagioclase contains 88–95mol% An. In addition, (4.0–5.9wt.%), Fe2O3 (6.9–10.0wt.%), K2O (0.01–0.03wt.%), scapolite (Table 1) has been found in an epidote-rich aggregate TiO2 (0.47–0.57wt.%), Zr (27–32ppm), and Nb (0.83– in a pillow core. 1.31ppm), but high CaO (17.4–20.8wt.%) and SiO2 (49.2– A. Polat et al. / Gondwana Research 11 (2007) 69–91 77

Table 2 TiO2, Ti/Zr, and Zr/Y ratios in the outer cores are comparable to Measured and recommended values for the USGS standards BHVO-1 and those in the inner cores (Table 3). Similarly, the chondrite- and BHVO-2 primitive mantle-normalized patterns are comparable to those of Element BHVO-2 BHVO-2 BHVO-1 BHVO-1 the inner cores (Fig. 7). (n=15) (n=10) The pillow rims are compositionally uniform at 8.1– Measured Recommended Measured Recommended 10.3wt.% MgO, 47.4–51.1wt.% SiO2, 11.7–13.2wt.% CaO, Li 4.5 5.0 4.63 0.53–0.57wt.% TiO2, 10.7–12.9wt.% Fe2O3, 592–670ppm V 331 317 398.1 Cr, and Mg-numbers (59–62) (Table 3). The ratios of Al2O3/ Cr 286 280 340.3 TiO2 (22–27) are super-chondritic, but Ti/Zr (96–102) ratios Co 45 45 44.55 – Ni 116 119 148.69 are sub-chondritic. Zr/Y (2.2 2.7) ratios range from sub- Cu 135 127 143.11 chondritic to super-chondritic (see Sun and McDonough, Zn 102 103 170.43 1989). In addition, they have the following trace element Rb 9.0 9.8 9.12 features: (1) depleted to slightly enriched LREE patterns (La/ Sr 390 389 390.5 Sm =1.01–1.39; La/Yb =0.75–1.23); (2) near-flat HREE Y 24.1 26.0 23.2 28 cn cn patterns (Gd/Ybcn =0.96–1.13); (3) negative Nb (Nb/ Zr 165 172 163.8 179 ⁎ ⁎ Nb 15.29 18.00 15.3 19 Nb =0.28–0.35) and Ti (Ti/Ti =0.71–0.92) anomalies; and Cs 0.10 0.10 0.13 (4) negative to positive Zr anomalies (Zr/Zr⁎=0.86–1.14) Ba 131 130 127.1 139 (Fig. 7). La 14.93 15 14.7 16 Ce 37.24 38 37.2 39 Pr 5.27 5.17 5.4 6.2. Gabbros and diorites Nd 24.08 25.0 23.6 25 Sm 5.95 6.20 5.99 6.4 Gabbros are characterized by 49–51wt.% SiO2, 0.51– Eu 2.00 1.97 2.06 0.90wt.% TiO2, 8.1–14.1wt.% MgO, 13.6–16.0wt.% Al2O3, Gd 6.17 6.30 6.21 6.4 and 9.0–13.2wt.% Fe2O3. There are large variations in Ni Tb 0.92 0.90 0.90 0.96 – – Dy 5.19 5.13 5.2 (121 234ppm) and Cr (246 1060ppm), and moderate varia- Ho 0.95 1.04 0.94 0.99 tions in Co (52–61ppm), V (186–264), Zr (28–44ppm), Y Er 2.54 2.49 (14–20ppm), and REE concentrations (e.g. La=2.2–8.0ppm) Tm 0.33 0.32 0.33 (Table 4). The ratios of Al2O3/TiO2 (16–27), Zr/Y (2.1–2.7), Yb 1.94 2.00 1.91 2.00 and Ti/Zr (94–121) range from sub-chondritic to super- Lu 0.27 0.28 0.27 0.29 Hf 4.39 4.10 3.91 4.4 chondritic values (see Sun and McDonough, 1989). In Ta 1.01 1.40 0.88 1.2 addition, they have the following trace element characteristics: Pb 1.77 1.85 2.6 (1) flat to slightly enriched REE patterns (La/Smcn =1.3–2.3, Th 1.47 1.20 1.36 1.1 Gd/Ybcn =1.0–1.4; 485428 is an outlier); and (2) large U 0.36 0.27 negative Nb (Nb/Nb⁎=0.29–0.50) and Ti (Ti/Ti⁎=0.60– 0.90) anomalies (Fig. 8; Table 4). Diorites have higher SiO2, but lower MgO, Fe2O3, Ni and Cr concentrations than those in 55.7wt.%) contents (Table 3) compared to modern average mid- gabbros (Table 4). Chondrite-and primitive mantle-normalized ocean ridge basalt (MORB) (see Hofmann, 1988). Mg-numbers trace element patterns of diorites are similar to those of gabbros range from 46 to 57 (Table 3). They display consistently higher (Fig. 8). Ni (267–403ppm), Cr (546–735ppm) and Co (41–70ppm) concentrations than average MORB. Al2O3/TiO2 ratios (26–28 6.3. Ultramafic pillows and cumulates versus 22) are slightly super-chondritic, whereas Ti/Zr (97–104 versus 115) ratios are slightly sub-chondritic. Zr/Y (2.0–2.8 Ultramafic pillows are compositionally variable at 10.1– versus 2.4) ratios range from sub-chondritic to super-chondritic 18.6wt.% MgO, 47.9–52.1wt.% SiO2, 10.7–12.5wt.% Fe2O3, (see Sun and McDonough, 1989). On chondrite- and primitive 330–660ppm Ni, 1300–5700ppm Cr, 62–93ppm Co, and mantle-normalized diagrams, they have the following trace Mg-number=64–77 (Table 5). They possess moderately element characteristics: (1) moderately fractionated LREE variable Al2O3 (8.7–11.3wt.%), TiO2 (0.37–0.51wt.%), patterns (La/Smcn =1.38–1.90; La/Ybcn =1.09–1.97), near flat CaO (8.5–12.7wt.%), Cr (1300–5700ppm), Ni (301– HREE patterns (Gd/Ybcn =0.92–1.13); and (3) variably nega- 659ppm) and V (136–199ppm) abundances (Table 5). Ti/Zr tive Nb (Nb/Nb⁎=0.27–0.39), Ti (Ti/Ti⁎=0.76–0.87), and Zr (76–116) and Zr/Y (2.03–2.77) ratios range from sub- ⁎ (Zr/Zr =0.78–1.00) anomalies (Fig. 7). chondritic to super-chondritic. Al2O3/TiO2 (24–25) ratios The outer pillow cores have higher MgO (5.3–8.4wt.%), are slightly super-chondritic. In addition, they have the SiO2 (51.3–58.4wt.%), K2O (0.10–0.18wt.%), and Na2O following geochemical characteristics: (1) LREE-enriched (0.81–2.5wt.%), but lower CaO (11.4–14.2wt.%) contents patterns (La/Smcn =1.24–2.31; La/Ybcn =1.18–2.06) and (2) than their inner counterparts (Table 3). Mg-numbers are negative Nb (Nb/Nb⁎=0.26–0.39), Zr (Zr/Zr⁎=0.81–0.98), consistently higher in the outer than inner cores (56–65 versus and Ti (Ti/Ti⁎=0.66–0.84) anomalies (Fig. 8). In comparison 46–57). Al2O3,TiO2, Cr, Ni, Co, Zr, and V contents, and Al2O3/ to pillows, cumulates have higher MgO (22.3–23.5wt.%), Ni 78 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Table 3 Major (wt.%) and trace (ppm) element concentrations and significant element ratios for mineralogically zoned pillow basalts Inner pillow cores 485441 485443 485444 485445 485446 485447 485448

SiO2 50.45 53.53 49.16 55.21 50.07 55.72 51.97 TiO2 0.57 0.47 0.55 0.53 0.52 0.51 0.47 Al2O3 15.36 13.01 14.66 13.89 14.61 13.59 12.39 Fe2O3 7.87 8.45 8.68 7.25 9.38 6.91 9.96 MnO 0.21 0.22 0.22 0.19 0.21 0.19 0.24 MgO 4.39 4.51 5.89 4.70 4.29 4.02 4.25 CaO 20.76 19.26 19.74 17.40 20.45 18.69 20.52 K2O 0.01 0.02 0.01 0.03 0.01 0.01 0.01 Na2O 0.34 0.48 1.07 0.75 0.41 0.31 0.15 P2O5 0.04 0.05 0.03 0.03 0.05 0.04 0.04 LOI 1.24 0.72 0.30 0.30 1.02 1.77 0.81 Mg-number 52 51 57 56 48 54 46

Cr 735 582 665 605 627 645 546 Co 63 54 64 70 52 41 56 Ni 403 369 366 383 329 267 294 Rb 4.3 2.5 0.9 1.1 4.5 1.3 1.4 Sr 80.4 71.6 100.9 85.6 111.6 73.8 117.5 Cs 0.8 0.0 0.1 0.1 0.3 1.3 0.5 Ba 25.7 15.2 27.9 26.7 35.1 7.5 9.5 Sc 47 37 46 44 43 43 38 V 233 223 186 193 221 214 232 Ta 0.09 0.06 0.08 0.09 0.08 0.09 0.07 Nb 1.31 0.83 1.21 1.15 1.07 1.07 1.19 Zr 32.5 27.3 30.0 31.6 29.3 31.4 27.1 Hf 1.14 0.96 1.06 1.10 1.05 1.11 0.92 Th 0.53 0.35 0.73 0.66 0.65 0.66 0.60 U 0.19 0.09 0.15 0.14 0.56 0.16 0.16 Y 12.6 12.0 13.3 12.3 11.2 11.4 13.3

La 2.82 3.40 3.07 3.22 2.89 2.79 2.40 Ce 6.72 6.85 7.24 7.11 6.58 6.36 5.77 Pr 0.92 0.79 0.99 0.97 0.89 0.87 0.81 Nd 4.47 4.03 4.80 4.51 4.23 4.04 3.82 Sm 1.46 1.28 1.50 1.34 1.33 1.20 1.22 Eu 0.47 0.51 0.57 0.54 0.51 0.50 0.47 Gd 1.92 1.64 2.11 1.92 1.84 1.73 1.76 Tb 0.35 0.31 0.37 0.35 0.32 0.32 0.33 Dy 2.45 2.00 2.59 2.42 2.21 2.22 2.34 Ho 0.53 0.43 0.55 0.50 0.48 0.46 0.49 Er 1.58 1.32 1.61 1.54 1.35 1.36 1.59 Tm 0.23 0.20 0.24 0.22 0.20 0.21 0.24 Yb 1.52 1.23 1.56 1.50 1.34 1.35 1.59 Lu 0.22 0.18 0.24 0.22 0.21 0.20 0.24

Cu 133 n.d. 320 172 108 118 43 Zn 65 41 69 54 55 47 62 Mo 0.6 n.d. 0.5 0.3 0.5 0.3 0.6 Ga 45 17 37 41 50 40 41 Pb 6.48 6.22 6.03 4.00 11.78 4.21 7.25 Li 42.37 n.d. 60.75 44.20 42.41 26.65 36.55 Bi 1.37 n.d. 0.43 0.47 2.74 0.24 2.57

La/Ybcn 1.34 1.97 1.41 1.54 1.54 1.48 1.09 La/Smcn 1.38 1.90 1.47 1.72 1.56 1.67 1.41 Gd/Ybcn 1.05 1.10 1.12 1.06 1.13 1.06 0.92 Eu/Eu⁎ 0.86 1.08 0.98 1.03 1.00 1.05 0.98 Ce/Ce⁎ 1.02 1.03 1.02 0.99 1.01 1.00 1.01 AL2O3/TiO2 26.9 27.8 26.7 26.0 28.1 26.7 26.4 Zr/Y 2.6 2.3 2.3 2.6 2.6 2.8 2.0 Ti/Zr 105 103 109 101 106 97 104 Nb/Nb⁎ 0.39 0.27 0.29 0.29 0.28 0.29 0.36 Zr/Zr⁎ 0.89 0.84 0.78 0.89 0.86 1.00 0.87 Ti/Ti⁎ 0.84 0.79 0.76 0.82 0.82 0.87 0.79 North 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ West 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ A. Polat et al. / Gondwana Research 11 (2007) 69–91 79

Outer pillow cores 485449485441 485450 485451485443 485453485444 485454 485445 485455 485446 58.27 58.20 51.33 51.39 56.61 58.42 0.57 0.52 0.56 0.58 0.54 0.48 15.25 13.72 14.02 14.57 14.89 14.36 6.25 8.78 10.20 10.83 8.19 6.22 0.17 0.16 0.22 0.20 0.16 0.17 5.34 5.58 7.86 8.24 5.88 5.79 12.01 11.52 14.16 11.41 11.68 13.60 0.11 0.10 0.11 0.18 0.18 0.11 1.99 1.37 1.50 2.56 1.82 0.81 0.03 0.04 0.04 0.02 0.05 0.03 0.74 0.68 3.76 0.80 0.56 0.49 63 56 60 60 59 65

649 603 634 688 620 584 38 45 52 63 58 39 175 182 253 319 257 212 2.7 2.3 2.2 1.8 2.2 2.9 76.2 65.3 137.9 74.0 86.5 89.2 0.0 0.0 0.0 0.0 0.0 0.4 49.7 30.8 39.0 48.7 53.3 39.8 46 42 47 46 43 39 226 206 243 251 225 208 0.09 0.08 0.10 0.09 0.08 0.08 1.21 1.17 1.29 1.22 1.32 1.07 30.7 29.3 32.8 34.1 32.6 29.0 1.09 0.93 1.13 1.18 1.02 0.95 0.70 0.67 0.76 0.69 0.70 0.65 0.19 0.17 0.25 0.26 0.36 0.24 14.0 12.7 13.7 14.4 13.0 12.4

3.09 2.39 3.21 2.64 4.34 3.08 7.42 6.17 7.40 6.72 10.45 6.92 1.07 0.86 1.03 0.98 1.23 0.93 4.79 4.16 4.74 4.60 5.36 4.35 1.49 1.27 1.34 1.43 1.43 1.38 0.52 0.46 0.50 0.52 0.56 0.48 2.09 1.92 2.06 2.18 1.97 1.84 0.38 0.33 0.35 0.38 0.37 0.33 2.53 2.30 2.48 2.60 2.47 2.29 0.54 0.49 0.54 0.57 0.51 0.49 1.68 1.50 1.64 1.68 1.57 1.46 0.24 0.22 0.25 0.25 0.22 0.21 1.64 1.51 1.56 1.62 1.51 1.45 0.25 0.21 0.25 0.25 0.22 0.21

119 32 51 5 23 9 88 82 112 123 117 88 0.3 0.2 0.3 0.1 0.3 0.2 40 36 40 37 46 43 4.57 4.37 6.64 3.26 4.49 3.97 43.79 38.16 45.71 70.04 39.46 34.89 0.15 0.07 0.09 0.06 0.09 0.06

1.35 1.14 1.48 1.16 2.06 1.52 1.48 1.34 1.72 1.32 2.19 1.60 1.05 1.05 1.09 1.11 1.08 1.05 0.90 0.90 0.92 0.90 1.02 0.92 1.00 1.05 1.00 1.02 1.11 1.00 26.6 26.6 25.0 25.0 27.5 29.9 2.2 2.3 2.4 2.4 2.5 2.4 111 105 102 102 100 98 0.30 0.33 0.30 0.33 0.27 0.27 0.80 0.89 0.91 0.93 0.82 0.83 0.80 0.81 0.83 0.81 0.79 0.74 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ Pillow rims (continued on next page) 485457 485458 485459 485460 485461 485462 485464 485465 485466 80 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Table 3 (continued) Pillow rims 485457 485458 485459 485460 485461 485462 485464 485465 485466

SiO2 49.48 48.52 51.97 47.44 47.47 50.32 51.12 48.22 49.89 TiO2 0.56 0.57 0.53 0.56 0.56 0.55 0.54 0.57 0.50 Al2O3 14.55 14.16 13.28 14.85 14.74 12.23 14.72 13.15 13.36 Fe2O3 11.57 12.09 11.19 12.70 12.85 12.34 10.65 12.82 11.82 MnO 0.23 0.22 0.22 0.24 0.24 0.24 0.19 0.26 0.21 MgO 8.65 9.45 8.59 9.38 9.40 10.32 8.21 10.40 9.80 CaO 12.89 13.18 12.08 12.98 12.90 11.83 11.71 13.06 12.35 K2O 0.15 0.16 0.15 0.19 0.21 0.25 0.18 0.15 0.16 Na2O 1.89 1.62 1.94 1.63 1.61 1.86 2.63 1.33 1.86 P2O5 0.04 0.04 0.04 0.03 0.03 0.04 0.05 0.03 0.05 LOI 0.82 0.89 0.95 1.00 0.93 1.00 0.76 1.01 0.87 Mg-number 60 61 60 59 59 62 60 62 62

Cr 670 642 592 631 623 604 631 640 603 Co 69 60 61 63 63 62 49 68 51 Ni 315 255 314 271 265 207 264 228 248 Rb 3.0 2.7 2.5 2.5 7.8 3.3 1.5 4.6 2.3 Sr 83.7 93.8 80.1 61.3 69.1 61.4 99.2 93.6 118.1 Cs 0.1 0.1 0.3 0.0 0.7 0.4 0.0 1.0 0.0 Ba 44.1 44.1 39.9 45.2 68.9 56.9 73.3 53.5 45.2 Sc 46 46 42 44 45 45 44 45 41 V 251 220 231 227 225 229 217 218 204 Ta 0.09 0.11 0.08 0.07 0.08 0.08 0.09 0.08 0.08 Nb 1.19 1.29 1.15 1.21 1.03 1.17 1.22 0.95 1.12 Zr 35.0 34.1 32.6 32.8 32.5 33.1 32.3 34.3 30.7 Hf 1.15 1.16 1.05 1.07 1.13 1.14 1.13 1.10 1.04 Th 0.64 0.70 0.63 0.62 0.62 0.59 0.80 0.89 0.66 U 0.19 0.17 0.17 0.16 0.09 0.13 0.15 0.14 0.15 Y 15.9 14.5 13.5 13.1 14.0 13.4 14.2 14.0 13.4

La 2.75 2.60 2.56 2.48 2.07 2.23 2.03 1.72 2.45 Ce 6.98 6.70 6.14 6.47 5.57 5.75 5.63 4.88 6.18 Pr 0.99 0.95 0.86 0.83 0.81 0.86 0.84 0.75 0.86 Nd 5.00 4.49 4.16 4.05 4.08 3.99 4.12 3.64 4.12 Sm 1.63 1.52 1.37 1.28 1.34 1.33 1.45 1.20 1.29 Eu 0.52 0.42 0.46 0.48 0.49 0.47 0.52 0.45 0.44 Gd 2.32 2.10 1.95 1.94 2.01 1.99 1.99 1.91 1.87 Tb 0.42 0.39 0.35 0.34 0.37 0.35 0.35 0.35 0.35 Dy 2.87 2.60 2.45 2.32 2.49 2.44 2.55 2.43 2.34 Ho 0.64 0.58 0.53 0.50 0.55 0.52 0.55 0.57 0.50 Er 1.89 1.70 1.56 1.50 1.57 1.53 1.72 1.67 1.49 Tm 0.28 0.24 0.23 0.21 0.24 0.23 0.24 0.24 0.22 Yb 1.83 1.53 1.50 1.50 1.65 1.50 1.57 1.64 1.47 Lu 0.27 0.23 0.23 0.22 0.24 0.23 0.24 0.25 0.23

Cu1142063249711 Zn 130 123 131 120 153 143 103 143 123 Mo 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 0.1 Ga 38 38 36 39 44 40 47 40 38 Pb 4.90 4.04 6.57 2.33 4.27 9.03 4.24 7.58 3.28 Li 50.27 46.07 53.58 62.46 70.83 63.28 35.33 38.25 41.25 Bi 0.14 0.06 0.11 0.06 0.09 0.10 0.05 0.10 0.07

La/Ybcn 1.08 1.21 1.23 1.19 0.90 1.07 0.93 0.75 1.19 La/Smcn 1.21 1.22 1.34 1.39 1.11 1.21 1.01 1.03 1.35 Gd/Ybcn 1.05 1.13 1.08 1.07 1.01 1.10 1.04 0.96 1.05 Eu/Eu⁎ 0.82 0.71 0.86 0.93 0.92 0.88 0.94 0.91 0.87 Ce/Ce⁎ 1.04 1.05 1.02 1.10 1.05 1.02 1.06 1.05 1.04 AL2O3/TiO2 26.0 24.8 25.2 26.4 26.3 22.2 27.3 23.2 26.7 Zr/Y 2.2 2.4 2.4 2.5 2.3 2.5 2.3 2.5 2.3 Ti/Zr 96 100 97 102 102 100 100 98 97 Nb/Nb⁎ 0.33 0.35 0.33 0.35 0.33 0.37 0.35 0.28 0.32 Zr/Zr⁎ 0.86 0.91 0.95 1.01 0.97 1.00 0.92 1.14 0.93 Ti/Ti⁎ 0.71 0.78 0.79 0.88 0.84 0.83 0.78 0.92 0.79 North 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ 64°44.906′ West 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ 49°51.827′ A. Polat et al. / Gondwana Research 11 (2007) 69–91 81

Fig. 7. Chondrite-normalized REE and primitive mantle-normalized trace element patterns for the inner, outer pillow cores and rims. Chondrite normalization values are from Sun and McDonough (1989) and primitive mantle normalization values are from Hofmann (1988).

207 206 (735–897ppm) and Mg-numbers (82–83), but lower TiO2 analysed grains gave Pb– Pb ages between 3082 and (0.26–0.28wt.%), Zr (10.5–15.6ppm), Y (7.1–8.4ppm), Cr 2998Ma, interpreted as detrital ages and indicative of (1575–1690ppm). Trace element systematics of cumulates are deposition at <2998Ma from a Mesoarchean source (Fig. 9, comparable to those of pillows (Fig. 8). Table 6). This is consistent with existing published ages for supracrustal rocks of the Ivisaartoq greenstone belt (e.g., Friend 6.4. Nd isotopes and Nutman, 2005a).

Seven samples from ultramafic pillows and cumulates have 8. Discussion been analysed for Sm–Nd isotope systematics (Table 5). They are characterized by variably positive initial εNd (3075Ma) 8.1. Origin of ocelli values (+1.29 to +4.97). As a group, cumulates have larger initial εNd values than pillows (+4.23 to +4.97 versus +1.29 to Ocelli are the felsic spherical (ellipsoidal in deformed rocks) +3.10). textural features, ranging mainly from 1 to 10cm in diameter. They occur in Archean mafic to ultramafic volcanic rocks (de 7. Geochronology results Wit et al., 1992; Kusky and Vearncombe, 1997; Dann, 2000; Anhaeusser, 2001). Ocelli occurring in the Ivisaartoq pillow Sample 493026 of a metasedimentary (volcanocalstic?) rock basalts are geometrically and mineralogically similar to those yielded only 13 zircons–most broken grains–that were 30– found in other Archean belts (see Gélinas et al., 1976; de Wit et 100μm diameter with aspect ratios of 1:2 to 1:2.5. Of these only al., 1992). To the best knowledge of the authors, ocelli in the 5 were large enough for laser ablation analyses. Back-scattered Ivisaartoq gabbros represent the first known Archean examples. electron imaging revealed a weakly defined oscillatory zonation Ocelli in pillow basalts have been cut by concentric cooling in some grains, though most were homogeneous. The five fractures and replaced by the stage I metasomatic mineral 82 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Table 4 Major (wt.%) and trace (ppm) element concentrations and significant element ratios for gabbros and diorites Gabbros Diorites 485428 485438 485467 485472 485476 485477 485485 485483 485484

SiO2 50.16 49.32 49.24 50.25 49.11 50.98 50.30 57.08 55.88 TiO2 0.90 0.64 0.51 0.59 0.61 0.62 0.62 1.14 0.90 Al2O3 14.64 15.24 13.58 14.99 15.32 16.04 15.88 16.86 15.96 Fe2O3 13.21 9.02 12.18 10.86 11.44 9.76 10.62 6.31 7.40 MnO 0.19 0.20 0.20 0.19 0.18 0.16 0.16 0.12 0.12 MgO 7.80 14.11 9.72 9.64 9.11 8.08 9.45 4.17 3.82 CaO 11.74 8.32 12.52 10.37 12.23 11.87 11.46 9.23 12.02

K2O 0.09 0.70 0.17 0.31 0.06 0.04 0.09 0.36 0.28 Na2O 1.21 2.44 1.83 2.74 1.88 2.40 1.38 4.69 3.53 P2O5 0.06 0.03 0.05 0.05 0.06 0.04 0.05 0.06 0.08 LOI 0.76 1.62 0.79 1.02 0.80 0.72 0.95 0.55 0.30 Mg-number 54 76 61 64 61 62 64 57 51 Cr 544 1061 332 375 292 323 246 187 189 Co 55 52 61 53 53 53 53 41 48 Ni 158 144 234 134 121 127 150 98 122 Rb 1.8 25.0 2.3 69.0 2.1 1.7 1.3 3.2 5.1 Sr 84 87 122 118 55 54 73 219 235 Cs 0.09 5.55 0.02 7.20 0.18 0.09 0.40 0.11 0.15 Ba 20.3 120.6 45.8 57.9 17.4 12.1 14.3 142.0 143.6 Sc 41 34 41 36 36 37 33 44 35 V 264 186 221 196 203 213 194 302 247 Ta 0.11 0.12 0.08 0.10 0.11 0.10 0.10 0.19 0.18 Nb 1.77 1.65 1.10 1.63 1.63 1.68 1.73 3.06 2.44 Zr 44.3 35.6 28.2 37.4 37.3 33.2 38.8 68.4 52.2 Th 0.217 0.532 0.726 0.465 0.628 0.503 0.503 1.161 0.942 U 0.035 0.100 0.160 0.149 0.176 0.257 0.151 0.355 0.298 Y 19.8 15.3 13.6 14.1 14.7 13.9 14.3 22.2 15.3 La 2.67 7.98 2.21 2.87 3.49 3.40 3.32 6.63 4.58 Ce 7.44 16.77 5.78 7.23 8.22 8.43 8.11 15.85 10.99 Pr 1.21 2.21 0.83 1.05 1.15 1.22 1.18 2.22 1.57 Nd 6.33 9.47 4.12 5.10 5.55 5.57 5.53 10.27 7.36 Sm 2.15 2.47 1.26 1.51 1.74 1.56 1.65 2.76 2.34 Eu 0.78 0.75 0.43 0.44 0.61 0.60 0.58 0.91 0.86 Gd 2.94 2.87 1.83 2.20 2.26 2.22 2.21 3.47 2.74 Tb 0.54 0.45 0.34 0.39 0.38 0.36 0.39 0.58 0.47 Dy 3.67 2.99 2.39 2.55 2.62 2.42 2.53 3.94 3.05 Ho 0.79 0.62 0.52 0.52 0.55 0.52 0.54 0.82 0.64 Er 2.44 1.85 1.56 1.58 1.68 1.53 1.55 2.50 1.93 Tm 0.36 0.27 0.23 0.22 0.24 0.22 0.23 0.37 0.31 Yb 2.41 1.74 1.51 1.53 1.56 1.49 1.54 2.58 1.90 Lu 0.35 0.27 0.21 0.22 0.23 0.23 0.22 0.38 0.29 Cu 169.2 15.2 8.6 27.7 72.5 44.4 100.3 71.4 28.5 Zn 102.2 187.2 119.3 102.9 117.2 116.4 91.4 162.4 117.7 Mo 0.11 0.31 0.12 0.12 0.18 0.23 0.12 0.33 0.57 Ga 38.7 50.3 39.3 41.3 40.5 37.9 36.5 62.8 58.6 Pb 1.0 6.0 3.5 4.4 6.4 7.5 4.2 5.4 3.8 Li 20.2 75.8 39.1 138.1 63.1 48.1 72.7 48.1 79.4 Bi 0.02 0.08 0.06 0.37 0.08 0.06 0.14 0.27 0.20

La/Ybcn 0.79 3.29 1.05 1.35 1.60 1.64 1.55 1.84 1.73 La/Smcn 0.89 2.32 1.26 1.36 1.44 1.56 1.44 1.72 1.40 Gd/Ybcn 1.01 1.37 1.00 1.19 1.20 1.23 1.19 1.11 1.20 Eu/Eu⁎ 0.94 0.86 0.87 0.74 0.94 0.99 0.92 0.89 1.04 Ce/Ce⁎ 1.01 0.98 1.05 1.02 1.01 1.02 1.01 1.01 1.01

AL2O3/TiO2 16 24 27 25 25 26 26 15 18 Zr/Y 2.24 2.32 2.07 2.66 2.53 2.38 2.72 3.08 3.40 Ti/Zr 121 106 108 94 99 112 95 98 104 Nb/Nb⁎ 0.84 0.29 0.31 0.51 0.40 0.47 0.48 0.40 0.43 Zr/Zr⁎ 0.84 0.51 0.87 0.94 0.84 0.79 0.90 0.90 0.88 Ti/Ti⁎ 0.88 0.59 0.83 0.79 0.76 0.82 0.79 0.90 0.88 North 64°44.464′ 64°44.962′ 64°44.941′ 64°44.767′ 64°44.815′ 64°44.815′ 64°44.448′ 64°44.450′ 64°44.452′ West 49°56.311′ 49°53.313′ 49°51.831′ 49°51.576′ 49°51.046′ 49°51.046′ 49°54.379′ 49°53.799′ 49°53.789′ A. Polat et al. / Gondwana Research 11 (2007) 69–91 83

Fig. 8. Chondrite-normalized REE and primitive mantle-normalized trace element patterns for gabbros, diorites, ultramafic pillows, and cumulates. Chondrite normalization values are from Sun and McDonough (1989) and primitive mantle normalization values are from Hofmann (1988). assemblage, suggesting that their formation predates the thermal alteration (Kusky and Polat, 1999; Polat and Kerrich, Mesoarchean seafloor hydrothermal alteration. Following 2006 and references therein). Despite the polyphase deforma- previous studies (Gélinas et al., 1976; Cawthorn et al., 1979; tion, several recent studies have provided a number of lines of Coltortietal.,1987), we interpret the felsic ocelli and physical and chemical evidence for seafloor hydrothermal surrounding mafic matrix in the Ivisaartoq rocks as solidified alteration in least deformed domains of Paleo- to Neoarchean immiscible liquids. greenstone belts (Polat et al., 2003; Terabayashi et al., 2003; Yamamoto et al., 2004; Weiershäuser and Spooner, 2005). 8.2. Metasomatic alteration These studies documented some similarities and differences between modern and Archean seafloor hydrothermal alteration Many Archean greenstone belts exhibit a complex history of processes. post-magmatic deformation that obliterated the primary mag- Field and microscopic observations suggest that the matic stratigraphy, volcanic structures, and more subtle Ivisaartoq greenstone belt has undergone at least two stages geochemical signatures of processes such as seafloor hydro- of calc-silicate metasomatic alteration. The stage I alteration is 84 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Table 5 Major (wt.%) and trace (ppm) element data for ultramafic pillows and cumulates Pillows Cumulates 485417 485418 485420 485481 485482 485473 485474 485475 485478

SiO2 47.94 48.41 52.06 50.39 49.75 50.34 48.73 48.84 48.16 TiO2 0.45 0.43 0.51 0.35 0.37 0.27 0.28 0.27 0.26 Al2O3 11.26 10.87 12.01 8.69 9.02 6.36 7.05 6.28 7.19 Fe2O3 11.67 12.51 11.17 11.08 10.65 9.21 10.42 9.90 10.32 MnO 0.18 0.22 0.19 0.20 0.18 0.17 0.17 0.17 0.18 MgO 14.81 14.90 10.12 18.61 17.10 23.09 23.51 22.29 23.19 CaO 12.33 11.26 12.66 8.46 11.50 10.23 9.56 12.03 10.39

K2O 0.14 0.15 0.01 1.33 0.03 0.02 0.01 0.01 0.01 Na2O 1.20 1.23 1.23 0.86 1.38 0.30 0.25 0.20 0.26 P2O5 0.03 0.02 0.04 0.03 0.02 0.01 0.02 0.01 0.03 LOI 1.44 1.50 0.82 2.05 1.80 4.15 4.64 3.97 4.66 Mg-number 72 70 64 77 76 83 82 82 82 Cr 5569 5713 2999 1298 1414 1575 1670 1654 1692 Co 73 93 62 81 81 84 84 83 84 Ni 466 605 331 608 659 845 852 735 897 Rb 14.2 12.3 0.4 275.8 2.3 0.8 0.7 0.7 0.6 Sr 101 43 90 69 76 33 31 33 38 Cs 1.37 4.67 0.13 60.84 0.53 0.23 0.20 0.13 0.16 Ba 26.4 45.0 29.7 216.1 9.6 1.9 10.9 6.1 2.6 Sc 36 36 37 28 26 24 24 26 23 V 178 178 199 142 136 104 112 107 107 Ta 0.07 0.07 0.09 0.08 0.09 0.05 0.05 0.06 0.05 Nb 0.92 0.94 1.26 1.11 1.29 0.77 0.73 0.69 0.70 Zr 28.4 22.3 33.3 25.5 29.5 12.5 14.6 15.6 10.5 Th 0.372 0.402 0.708 0.683 0.825 0.339 0.259 0.356 0.238 U 0.126 0.101 0.192 0.148 0.286 0.293 0.668 0.183 0.284 Y 10.4 11.0 12.3 10.9 10.7 7.0 7.1 8.4 7.4 La 1.95 2.25 3.31 3.56 3.34 1.69 1.59 1.82 1.92 Ce 4.92 4.95 7.83 7.80 7.96 4.10 3.75 4.26 4.14 Pr 0.73 0.76 1.06 1.00 1.04 0.52 0.54 0.63 0.56 Nd 3.59 3.66 4.96 4.34 4.57 2.35 2.43 2.89 2.55 Sm 1.13 1.11 1.39 1.10 1.17 0.73 0.70 0.87 0.73 Eu 0.37 0.44 0.46 0.41 0.41 0.23 0.22 0.21 0.23 Gd 1.52 1.64 1.85 1.54 1.54 0.98 0.98 1.22 1.06 Tb 0.27 0.29 0.33 0.25 0.27 0.18 0.17 0.22 0.17 Dy 1.90 1.96 2.28 1.84 1.75 1.20 1.26 1.52 1.24 Ho 0.40 0.42 0.48 0.38 0.39 0.26 0.27 0.33 0.26 Er 1.23 1.30 1.46 1.19 1.15 0.85 0.82 0.99 0.83 Tm 0.18 0.19 0.21 0.18 0.18 0.13 0.12 0.14 0.12 Yb 1.19 1.24 1.40 1.24 1.21 0.84 0.81 1.03 0.86 Lu 0.18 0.18 0.21 0.18 0.19 0.13 0.12 0.15 0.12 Cu 58.6 15.5 78.3 2.6 5.2 10.8 14.4 6.7 19.1 Zn 82.9 93.5 108.7 92.3 78.3 79.2 91.2 80.6 85.5 Mo 0.11 0.07 0.19 2.57 0.08 0.18 0.27 0.11 0.30 Ga 27.6 30.7 32.2 52.6 23.9 15.0 17.8 16.1 16.9 Pb 0.6 2.2 1.7 1.9 1.3 1.6 1.7 1.8 0.3 Li 104.9 129.0 32.2 149.1 74.5 5.8 4.4 9.5 7.9 Bi 2.46 0.67 0.07 0.64 1.03 0.18 0.33 0.50 0.67

La/Ybcn 1.18 1.30 1.70 2.06 1.98 1.45 1.41 1.27 1.60 La/Smcn 1.24 1.45 1.70 2.31 2.05 1.65 1.62 1.50 1.89 Gd/Ybcn 1.06 1.09 1.09 1.03 1.05 0.97 1.00 0.98 1.02 Eu/Eu⁎ 0.86 0.99 0.88 0.96 0.93 0.82 0.81 0.61 0.78 Ce/Ce⁎ 1.01 0.93 1.02 1.01 1.05 1.07 1.00 0.97 0.98

AL2O3/TiO2 25 25 24 25 24 24 25 23 27 Zr/Y 2.72 2.03 2.70 2.34 2.77 1.78 2.07 1.86 1.42 Ti/Zr 94 116 92 81 76 128 115 103 150 Nb/Nb⁎ 0.39 0.36 0.30 0.26 0.28 0.37 0.41 0.31 0.37 Zr/Zr⁎ 0.98 0.77 0.88 0.81 0.88 0.66 0.78 0.68 0.53 Ti/Ti⁎ 0.84 0.79 0.78 0.66 0.68 0.78 0.83 0.64 0.74 147Sm/144Nd 0.1853 0.1732 0.1618 0.1584 0.1795 0.1787 0.1764 143Nd/144Nd 0.512549 0.512226 0.512086 0.511967 0.512504 0.512512 0.512479 A. Polat et al. / Gondwana Research 11 (2007) 69–91 85

Table 5 (continued) Pillows Cumulates 485417 485418 485420 485481 485482 485473 485474 485475 485478

εNd (at 3075Ma) 2.81 1.29 3.10 2.10 4.23 4.72 4.97 North 64°44.363′ 64°44.375′ 64°44.334′ 64°44.515′ 64°44.515′ 64°44.763′ 64°44.763′ 64°44.763′ 64°44.761′ West 49°56.124′ 49°56.130′ 49°56.183′ 49°53.299′ 49°53.361′ 49°51.318′ 49°51.318′ 49°51.318′ 49°51.332′ characterized by the formation of an epidote (now mainly Significant silicification, with subordinate epidotization, of diopside)+quartz±plagioclase±amphibole (hornblende) as- pillow basalts during hydrothermal alteration of modern oceanic semblage in the inner pillow cores, and pillow interstitials. In crust, has also been documented from mid-ocean ridges the pillow cores, alteration occurs primarily around the drainage (Humphris and Tivey, 2000; Teagle and Alt, 2004;and cavities. This assemblage also occurs along the pillow-gabbro references therein). An extensive occurrence of quartz-filled contacts and fractures within gabbros. Collectively, the stage I pillow drainage cavities, interstitials, vesicles, and veins has metasomatic assemblage appearstohaveformedduring recently been reported from the Neoarchean Abitibi greenstone intrusion of gabbroic to dioritic dykes, predating regional belt, Canada (Weiershäuser and Spooner, 2005). The chemical metamorphism and deformation. Similar spatial and temporal composition of fluid inclusions is consistent with the precip- relations between late-stage plagiogranite and quartz diorite itation of quartz grains in these rocks during Neoarchean sea- intrusions and the formation of epidosites and mineralized floor hydrothermal alteration (Weiershäuser and Spooner, zones in the sheeted dike complex and the extrusive sequence 2005). Although we do not have chemical evidence from have been documented from the Semail ophiolite in Oman fluid inclusions, on the basis of field relationships, and (Stakes and Taylor, 2003). These intrusions appear to have comparing with the Abitibi quartz formations, we infer that provided both heat and hydrothermal fluids necessary for the the quartz filling the drainage cavities and interstitials in the alteration of the host upper crustal rocks and for the replacement Ivisaartoq pillow basalts likely precipitated during Mesoarchean of the pre-existing igneous minerals (pyroxene, ilmenite, and seafloor hydrothermal alteration. Some pillow rims in the calcic plagioclase) in these rocks by alteration minerals Ivisaartoq greenstone belt underwent extensive silica alteration. (chlorite, amphibole, albite, epidote, and quartz). Copper In some locations pillows have also been completely silicified enrichments in the lavas and dikes above the plagiogranite (Fig. 4d). Given the high MgO values in the non-silicified rims, bodies in Oman suggest that these intrusions had a major role in we suggest that silicification postdates chloritization of the rims the redistribution of heavy metals and precipitation of massive (cf. Humphris and Tivey, 2000 and references therein). sulphide deposits in the ophiolite (Stakes and Taylor, 2003). The stage II metasomatic assemblage occurs as calc-silicate Similarly, we suggest that the gabbroic intrusions in the veins and boudins that are concordant to discordant to the dominant extrusive rocks in the Ivisaartoq greenstone belt were foliation planes in the replaced host rocks, consistent with a pre-to responsible for widespread alteration and mineralization as syn-metamorphic origin (Fig. 3a–c). Most of these veins are documented here. The formation of the stage I metasomatic spatially associated with shear zones, which probably formed assemblage in drainage cavities, pillow interstitials, and during terrane accretion. The stage II metasomatic assemblage is intrusive contacts suggests therefore that these spaces acted as composed dominantly of granular diopside+garnet+hornblende conduits for the transportation of hydrothermal fluids. +plagioclase+quartz±vesuvianite±scapolite ±epidote±calcite ±scheelite. Diopside constitutes about 70% of the assemblage. Thin (1–3cm) veins of this assemblage crosscut the pillow rims and cores. The formation of the calc-silicate assemblage that contains a garnet+diopside pair is consistent with middle- to upper-amphibolite facies metamorphic conditions (Appel, 1997). Given that the stage II assemblage is spatially associated with the highly strained contact (shear zone) between the lower and upper amphibolite units, it is suggested that this assemblage was formed during a regional tecto- nothermal metamorphic event. The stage I metasomatic assemblage has been folded by an early phase of deformation, suggesting that they formed prior to isoclinal folding of the belt. However, the spatial and temporal relationships between the stage II metasomatism and isoclinal folding are not well constrained. In addition to the above metasomatic assem- blages, there are epidote-dominated veins or patches appear to replace diopside within both stage I and stage II assemblages. Fig. 9. Concordia plot of U–Pb zircon age data for detrital zircon grains from It appears that these epidote veins are associated with brittle metasedimentary rock sample 493026. deformation that might have accompanied retrograde 86 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Table 6 LA-SF-ICPMS U–Th–Pb data for detrital zircon grains from sample 493026 Analysis no. U Pb 206Pb/ Th/U 206Pb/ 1σ 207Pb/ 1σ rho 208Pb/ 1σ 207Pb/ 1σ 206Pb/ 207Pb/ Discordance, (ppm) (ppm) 204Pb 238U (%) 235U (%) 232U (%) 206Pb (%) 238U 206Pb % 493026-05 134 93 8745 0.66 0.5622 1.3 181.623 1.4 0.94 0.1413 0.64 0.2343 0.5 2876±76 3082±15 4.1 493026-07 116 75 12,159 0.59 0.523 1.8 160.389 1.9 0.95 0.152 1.9 0.2224 0.6 2712±100 2998±19 5.8 493026-08 60 44 14,205 0.8 0.5816 1.5 183.701 1.7 0.87 0.1365 0.62 0.2291 0.9 2955±88 3046±27 1.8 493026-10 52 41 21,922 0.83 0.6414 1.9 204.339 1.9 0.96 0.1208 2 0.231 0.5 3195±120 3059±16 −2.7 493026-11 202 135 0.42 0.5754 1.4 184.165 1.5 0.93 0.1444 0.44 0.2321 0.5 2930±80 3067±17 2.7 metamorphism during a post-orogenic phase of the tectonic Banerjee et al., 2000; Humphris and Tivey, 2000; Gillis and evolution of the Ivisaartoq region. Banerjee, 2000; Teagle and Alt, 2004). Convection of seawater through the oceanic crust leaches large quantities of sulphur, 8.3. Element mobility and siderophile and chalcophile elements, including Fe, Cu, Zn, and Pb; deposition of these elements as sulphides on the As described above, the Ivisaartoq pillow basalts are seafloor generates VMS and amorphous Fe-oxides and hydro- mineralogically and chemically zoned (Table 3). The rims xides (Seyfried et al., 1991; Alt et al., 1998; Alt, 1999; Alt and have higher contents of Fe2O3, MgO, MnO, and K2O, whereas Teagle, 2000; Humphris and Tivey, 2000). The experimental the inner and outer cores possess higher concentrations of CaO, results of Seyfried and Janecky (1985) suggest a strong increase and Na2O and SiO2, respectively, consistent with the mobility of in Fe concentrations in hydrothermal fluids with increasing these elements during post-magmatic alteration. Similarly, large temperature. Many supra-subduction zone ophiolites show variations in Ba, Sr, Pb, Rb, Cs, Li, U, Zn, and Cu contents evidence for high-temperature hydrothermal discharge in the indicate the mobility of these elements. Compared with the form of VMS and/or metalliferous sedimentary rocks (Robert- pillow cores, the rims have lower abundances of LREE and La/ son and Varnavas, 1993; Varga et al., 1999; Gillis and Banerjee, Smcn ratios, indicating the some loss of these elements. It is 2000; Banerjee and Gillis, 2001). The presence of pyrite within likely that LREE loss took place during chloritization of the quartz-filled drainage cavities and pillow interstitials suggest rims (cf. Humphris and Tivey, 2000; Teagle and Alt, 2004). The that they acted as channels for the transportation of Fe-rich mobility of Rb, K, Na, Sr, Ba, Fe, Mg, and Pb in Phanerozoic fluids. As a corollary, we propose that pyrite-rich felsic layers in and Archean pillow basalts is well documented, and this study the Ivisaartoq greenstone belt resulted from high-temperature confirms the mobility of these elements in the Ivisaartoq pillow alteration of the Mesoarchean oceanic crust. Alternatively, some basalts (see Condie et al., 1977; Kerrich and Fyfe, 1981; Polat of these felsic rocks may have formed as volcanoclastic rocks and Hofmann, 2003; Polat et al., 2003). Although REE are above hydrothermal discharge zones (see Friend and Nutman, generally regarded as immobile during hydrothermal alteration 2005a). and low-grade metamorphism, it has been shown that LREE can be mobilized during intense seafloor hydrothermal alteration 9. Conclusions and implications (Ludden and Thompson, 1979; Seifert et al., 1985; Valsami and Cann, 1992; Gillis et al., 1992; Ridley et al., 1994; Humphris et 9.1. Sea-floor hydrothermal alteration and formation of al., 1998; Wells et al., 1998; Teagle and Alt, 2004). In contrast epidosites to the above elements, Al2O3,TiO2, Th, Zr, Y, and Cr, Ni, Co, Ga, HREE display minor variations between the cores and rims, Thermal evolution models predict higher mantle tempera- suggesting that these elements were relatively immobile during tures in the Archean than at present (Fyfe, 1978; Bickle, 1986; seafloor hydrothermal alteration. Similarly, REE, HFSE (Ti, Galer and Mezger, 1998). High heat flow and intense magmatic Nb, Ta, Zr, Y) in gabbros, diorites, ultramafic pillow basalts, activity in the Archean would have promoted vigorous and cumulates display coherent primitive mantle-and chondrite- convection of seawater through the oceanic crust, both at normalized patterns, indicating that these elements were spreading centres and hot spots (Fryer et al., 1979). Therefore, relatively immobile during post-magmatic alteration (see also intense hydrothermal metasomatism would have substantially Ludden et al., 1982). hydrated and modified the primary chemical composition of the Archean oceanic crust. 8.4. Origin of pyrite-rich layers Given the petrographic observations indicating that epidote in pillow cores was replaced by diopside, we suggest that the Modern volcanogenic massive sulphides (VMS) are associ- stage I epidote+quartz±plagioclase±amphibole (chlorite?) ated with hydrothermal vents found in mid-ocean ridges, metasomatic assemblage formed as epidosites under greenschist forearcs, backarcs, and hot spots where seawater moves down facies metamorphic conditions during seafloor hydrothermal through the fractures in the oceanic crust, warms up to alteration. This early seafloor hydrothermal metamorphism was temperatures of over 400°C, interacts with oceanic crust, and later overprinted by middle- to upper-amphibolite facies then rises to the surface and discharges through hydrothermal regional metamorphism. During amphibolite-grade metamor- vents (Alt et al., 1998; Alt, 1999; Alt and Teagle, 2000; phism, epidosites underwent a major re-equilibration, resulting A. Polat et al. / Gondwana Research 11 (2007) 69–91 87 in the replacement of epidote by diopside. The precise ages of metasomatism reflects an interaction of pillow basalts with the stage I and stage II metasomatic events, and the gap between highly reactive, discharging fluids similar to those at modern the two are unknown. Intrusion of the belt by weakly deformed high-T °C hot springs on mid-ocean ridges (cf. Harper et al., 2963Ma granites (Friend and Nutman, 2005a and references 1988; Schifmann and Smith, 1988; Seyfried et al., 1988). therein) indicate that both metasomatic events formed shortly Despite the lithological similarities and the spatial and after the deposition of the belt between 3075 and 3000Ma (see temporal relationships between the alteration/mineralization Friend and Nutman, 2005a). zones and the late-stage plutonic intrusions as heat and fluid The presence of relict epidosites in the Ivisaartoq greenstone sources, there are several major differences in the characteristics belt has important implications for the geochemical modifica- of hydrothermal alteration and mineral assemblages between tion of the oceanic crust and geodynamic processes in the the basalts of the Ivisaartoq greenstone belt and those of the Mesoarchean. The principal implications are: (1) chemical modern oceanic crust and ophiolites. These include (1) presence variations in the Ivisaartoq pillow basalt rims and cores suggest of chlorite as a mafic mineral in Phanerozoic epidosites, a complex history of seafloor alteration and considerable whereas amphibole is the main mafic mineral in the Ivisaartoq redistribution of many elements; (2) given the observations that greenstone belt counterparts (although actinolite here may hydrothermal alteration of modern ocean-floor basalts are represent a late metamorphic mineral); (2) metasomatic largely controlled by sea-floor spreading processes, the alteration is zoned in many ophiolites, which is not clear in chemical reorganization of the Ivisaartoq pillow basalts may the Ivisaartoq greenstone belt; (3) the lack of reported diopside have resulted from Mesoarchean sea-floor spreading processes formation in modern ocean floor and ophiolitic pillow basalts in one of a variety of tectonic settings ranging from mid-ocean (also diopside, even in the pillow cores, most likely represents a ridges to back-arc, intra-arc, or fore-arc environments (Dilek et later formed regional metamorphic mineral, rather than a al., 1998; Dilek and Flower, 2003; Kusky, 2004); and (3) product of seafloor hydrothermal alteration (Note that diopside because Phanerozoic epidosites form in upflow zones at the replaces epidote); and (4) the presence of stockwork veins and base of ore-forming hydrothermal systems, their occurrence in alteration zone that can be traced into sheeted dyke complex the Ivisaartoq greenstone belt can explain the origin of the beneath VMS deposits in Phanerozoic ophiolites, in the pyrite-rich layers in the belt (cf. Teagle, 1993; Banerjee et al., Ivisaartoq greenstone belt these geological features have not 2000 and references therein). been recognized. It is plausible that alteration features widely The formation of epidosites in the more permeable pillow documented from the Phanerozoic ophiolites may have been cores and interstitials suggest that hydrothermal fluids may obliterated in the Ivisaartoq greenstone belt during multiple have channeled through these cavities. Heat and fluids derived episodes of intense deformation and metamorphism. from gabbroic and dioritic intrusions may have driven the hydrothermal convection, and pillow drainage and interstitial 9.2. Geodynamic setting and petrogenesis cavities likely have created permeability for fluid flow. Similar features have been described from the late Ordovician Pillow basalts, gabbros, diorites, and cumulates display Solund-Stavfjord ophiolite in western Norway (Furnes et al., similar trace element characteristics, suggesting similar 2001), from the late Cretaceous Troodos ophiolite in Cyprus sources and petrogenetic origin for all these spatially and (Eddy et al., 1998), and from the late Cretaceous Semail temporarily associated rock types. The slightly LREE- ophiolite in Oman (Stakes and Taylor, 2003). In each of these enriched and HFSE-depleted trace element patterns of the Phanerozoic examples the oceanic crust is heterogeneous in Ivisaartoq volcanic and intrusive rocks are consistent with a its internal architecture because of the late-stage intrusions of subduction zone geochemical signature (Saunders et al., 1991; gabbroic to plagiogranitic stocks and plutons into the pre- Hawkesworth et al., 1993; Pearce and Peate, 1995). Near-flat existing igneous complex, and the extensive hydrothermal HREE patterns are furthermore consistent with a shallow alteration in sheeted dikes and extrusive rocks has been melting depth. High MgO, Ni and Cr concentrations and Mg- observed to be associated with these intrusions, which acted numbers in ultramafic pillows and cumulates are consistent as shallow point sources of heat, fluids, and metals. Thus, we with a picritic composition. Tertiary island arc picrites have suggest that the mafic-ultramafic rocks preserved in the been documented in the Solomon and New Hebrides (Vanuatu 3075Ma Ivisaartoq greenstone belt display a hydrothermal arc) oceanic island arcs (Ramsay et al., 1984; Eggins, 1993; alteration history analogous to that of the Phanerozoic Schuth et al., 2004). In the Solomon Islands, picrites occur ophiolites and the modern oceanic crust. only in New Georgia Island above the subducting Woodlark Intense chemical, textural, and mineralogical changes in the spreading centre. In the New Hebrides subduction system, the Ivisaartoq lavas require the transportation of a large volume of overriding plate is currently undergoing extension east of the hydrothermal fluids through pillow basalts at high fluid/rock Vanuatu arc, forming a supra-subduction oceanic crust within ratios (cf. Richardson et al., 1987; Seyfried et al., 1988; Harper, the North Fiji Basin (Pearce, 2003). Given the geochemical 1995). The inner pillow cores are enriched in Ca and depleted in characteristics of the rock samples we analyzed in this study Mg and Na (Table 3), compared to non-metasomatized outer and the hydrothermal alteration features reminiscent of their cores. High MgO contents (average 9.36wt.%) in the rims may Phanerozoic counterparts, we suggest that the Ivisaartoq have resulted from chloritization during seafloor hydrothermal picrites and peridotites may have formed in a suprasubduction alteration (cf. Humphris and Tivey, 2000). This distinctive zone tectonic setting. 88 A. Polat et al. / Gondwana Research 11 (2007) 69–91

Given the presence of the primary textures and rare fresh cpx Tethyan ophiolites in the eastern Mediterranean region (i.e. in cumulates, the Nd isotopic composition (average εNd = Dilek and Flower, 2003). +4.64) in these rocks is likely to reflect the primary isotopic signature of the Mesoarchean mantle beneath the Ivisaartoq 9.3. Accretionary tectonics in the Nuuk region arc (sensu lato). Large positive initial εNd (+4.21 to +4.97) values are consistent with a strong Sm/Nd fractionation and a On the basis of detailed field observations, contrasting long-term depleted mantle source region. The depleted Nd ages, and metamorphic histories, it has been shown that the isotopic signatures and low LREE and HFSE (Nb, Ta, Zr, Ti, Nuuk region in southern West Greenland is a collage of Y) abundances indicate that the mantle source region had Paleo-to Neoarchean terranes, assembled in several accretion- experienced significant melt extraction prior to 3075Ma. The ary Archean tectonothermal events, involving accretion of concentrations of HFSE (e.g., Nb, Zr, Ti) in the Ivisaartoq exotic terranes (e.g., arcs, continental fragments) by horizontal volcanic rocks are much lower than those in modern N- tectonics (Bridgwater et al., 1974; McGregor et al., 1991; MORB, indicating that the Mesoarchean mantle beneath the Friend et al., 1996; Komiya et al., 1999; Garde et al., 2000; Ivisaartoq greenstone belt (sensu lato) was more depleted in Nutman et al., 2002, 2005; Friend and Nutman, 2005a,b; and these elements compared to the modern depleted upper references therein). Structural, geochronological, and trace mantle. The light-REE-enriched patterns of the Ivisaartoq element geochemical studies in the Paleoarchean Isua green- rocks, however, imply that the depleted mantle source of stone belt suggest that the ∼3800Ma island arc picrite-chert- these rocks must have been fertilized by some processes such BIF sequence and the ∼3700Ma boninite-chert-BIF sequence as metasomatism of the subarc mantle peridotite before or during were juxtaposed by accretionary Phanerozoic-like plate partial melting took place at about 3075Ma (cf. Polat and tectonic processes operating in the Mesoarchean (see Nutman Münker, 2004). Slab-derived hydrous fluids, with sub-chon- et al., 2002; Polat et al., 2002; Polat and Hofmann, 2003). dritic, Nb/La, Nb/Th, Zr/Sm, Sm/Nd, and Ti/Gd ratios, were Recently, Friend and Nutman (2005a) have shown that the likely the main cause of the mantle fertilization (metasomatism). Paleoarchean Isukasia (ca. 3800–3600Ma) and the Mesoarch- Interaction between hydrous fluids and the depleted mantle ean Kapisilik (ca. 3000Ma) terranes were tectonically wedge (with super-chondritic Sm/Nd ratio) may have played an juxtaposed at about 2960Ma (Fig. 1). This was followed by important role for the initiation of partial melting and for the the accretion of the Neoarchean (ca. 2800Ma) Tre Brødre generation of LREE-enriched, HFSE-depleted trace element terrane to the composite Isukasia-Kapisilik terrane. Subduc- patterns in the Ivisaartoq volcanic rocks. We suggest that high tion zone geochemical signatures and ophiolite-like crustal MgO, Ni, and Cr concentrations in these rocks reflect a large lithologies and alteration patterns in the Ivisaartoq greenstone degree of partial melting in mantle wedge during subduction of belt are consistent with the accretion of juvenile oceanic young, hot oceanic crust (see Polat and Kerrich, 2006). fragments to the pre-existing continental blocks (e.g. Isuakasia The Ivisaartoq greenstone belt is composed dominantly of terrane) in the Nuuk region through a forearc-continent pillow basalts and gabbroic sills and dykes, resembling the collision. Dilek and Ahmed (2003) and Stern et al. (2004) upper-most section of an idealized, Penrose-type ophiolitic have documented a similar evolutionary path for the oceanic crust. However, the preserved upper-crustal oceanic Proterozoic Arabian-Nubian Shield ophiolites via a series of crust in the Ivisaartoq greenstone belt is quite similar to the collisions of arc-trench complexes with continental margins. Jurassic Alpine-Apennine ophiolites, which have been defined Consequently, we think that the plate tectonic processes as remnants of a “Hess-type” ancient oceanic crust developed in responsible for the formation and emplacement of oceanic the slow-spreading Ligurian ocean (Dilek, 2003). In a Hess-type crust, as we know from the Phanerozoic, appear to have been ancient oceanic crust, pillow basalts, ophicalcites and chert are in operation during much of the Precambrian evolution of the seen to rest directly on serpentinized peridotites and gabbros, Earth. with no sheeted dike intrusions, and gabbroic dikes, stocks, and plutons are intrusive into the peridotites and locally into the Acknowledgements overlying pillow lavas (Dilek, 2003). The Hess-type oceanic crust is interpreted to form in magma-poor spreading environ- We thank A. Trenhaile and R. Kerrich for reviewing the ments, where periods of tectonic extension keeping pace with initial draft of the manuscript. Reviewers T. Kusky and H. the ongoing seafloor spreading attenuates the new oceanic crust Smithies are acknowledged for their constructive and incisive and exhumes the lower crustal and upper mantle rocks, comments that have resulted in significant improvements to the analogous to the tectonomagmatic evolution of slow-spreading paper. S. Maruyama and M. Santosh are acknowledged for oceanic crust at the mid-Atlantic Ridge (Cannot et al., 1995; inviting us to submit this contribution to Gondwana Research. Dilek et al., 1998; Dilek, 2003). Similar sequences are We thank A. Nutman, C. Friend, A. Garde, J. van Gool, H. documented from several other proposed Archean ophiolites Stendal, A. Steenfelt, D. Frei, O. Stecher, C. Münker, and N. (Kusky, 2004). While this interpretation is plausible for the Kelly for helpful discussions in the field. T. Bonli, J.C. igneous evolution of the Ivisaartoq greenstone belt “oceanic Barrette, J. Gagnon, and Z. Yang are acknowledged for their crust”, its inferred supra-subduction zone geochemical affinity help during geochemical analyses, and D. Frei for geochro- suggests to us that it may have likely formed in a Mesoarchean nology support. This is a contribution of NSERC grants forearc or juvenile island arc system, similar to some of the 250926 to AP and 83117 to B. Fryer. R. Frei is supported by A. Polat et al. / Gondwana Research 11 (2007) 69–91 89

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