Neoproterozoic arc–back-arc system in the Central Eastern Desert of Egypt: Evidence from supra-subduction zone ophiolites

Lithos 120 (2010) 293–308

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Lithos j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / l i t h o s

Neoproterozoic arc–back-arc system in the Central Eastern Desert of Egypt: Evidence from supra-subduction zone ophiolites E.S. Farahat ⁎ Department of Geology, Minia University, El-Minia 61519, Egypt

a r t i c l e

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Article history: Received 27 May 2010 Accepted 12 August 2010 Available online 20 August 2010 Keywords: Neoproterozoic ANS Ophiolites Boninite Arc–back-arc Egypt

a b s t r a c t Ophiolites are widely distributed in the Central Eastern Desert (CED) of Egypt, occurring as clusters in the northern (NCEDO) and southern (SCEDO) segments. Mineralogical and geochemical data on the volcanic sections of Wizer (WZO) and Abu Meriewa (AMO) ophiolites as representatives of the NCEDO and SCEDO, respectively, are presented. The WZO volcanic sequence comprises massive metavolcanics of MORB-like compositions intruded by minor boninitic dykes and thrust over island-arc metavolcanic blocks in the mélange matrix. Such transitional MORB-IAT-boninitic magmatic affinities for the WZO metavolcanics suggest that they most likely formed in a protoarc–forearc setting. Chemical compositions of primary clinopyroxene and Cr-spinel relicts from the WZO volcanic section further confirm this interpretation. The compositional variability in the WZO volcanic sequence is comparable with the associated mantle rocks that vary from slightly depleted harzburgites to highly depleted harzburgites containing small dunite bodies, which are residues after MORB, IAT and boninite melt formation, respectively. Source characteristics of the different lava groups from the WZO indicate generation via partial melting of a MORB source which was progressively depleted by melt extraction and variably enriched by subduction zone fluids. MORB-like magma may have been derived from ~ 20% partial melting of an undepleted lherzolite source, leaving slightly depleted harzburgite as a residuum. The generation of island-arc magma can be accounted for by partial melting (~15%) of the latter harzburgitic mantle source, whereas boninites may have been derived from partial melting (~20%) of a more refractory mantle source previously depleted by melt extraction of MORB and IAT melts, leaving ultra-refractory dunite bodies as residuum. The AMO volcanic unit occurs as highly deformed pillowed metavolcanic rocks in a mélange matrix. They can be categorized geochemically into LREE-depleted (La/YbCN = 0.41–0.50) and LREE-enriched (La/YbCN = 4.7– 4.9) lava types that show an island arc to MORB geochemical signature, respectively, signifying a back-arc basin setting. This is consistent, as well, with their mantle section. Source characteristics indicate depleted to slightly enriched mantle sources with overall slight subduction zone geochemical affinities as compared to the WZO. Generally, CED ophiolites show supra-subduction zone geochemical signature with prevalent island arc tholeiitic and minor boninitic affinities in the NCEDO and MORB/island-arc association in the SCEDO. Such differences in geochemical characteristics of the NCEDO and SCEDO, along with the abundance of mature island arc metavolcanics which are close in age (~750 Ma) to the ophiolitic rocks, general enrichment in HFSE of ophiolites from north to south, and lack of a crustal break and major shear zones, is best explained by a geotectonic model whereby the CED represents an arc–back-arc system above a southeast-dipping subduction zone. © 2010 Elsevier B.V. All rights reserved.

1. Introduction Ophiolites are key features for understanding the geologic history of orogenic belts, particularly in the Arabian–Nubian Shield (ANS) where they are widely distributed and range in ages from 870 to

⁎ Tel.: + 20 86 2378500; fax: + 20 86 2363011. E-mail address: [email protected]. 0024-4937/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2010.08.017

627 Ma (Dilek and Ahmed, 2003; Stern et al., 2004). Such remarkably broad abundance of ophiolites led Stern et al. (2004) to describe the ANS as a “massive graveyard of Neoproterozoic oceanic lithosphere.” The ANS is exposed around the Red Sea in the western Arabian Peninsula, southern Israel, Jordan and Sinai, Egyptian Eastern Desert, the Red Sea Hills of Sudan, Eritrea and as far south as Ethiopia and Yemen. It is an excellent example of juvenile continental crust formation in association with plate tectonics. The ANS is a part of the East African–Antarctic Orogen (Jacobs and Thomas, 2004) which is

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related to consolidation of East and West Gondwana. Lateral crustal growth through arc–arc accretion has been proposed as the main model for the evolution of the ANS (Kröner et al., 1987, 1991; Stern, 2002). However, possible contribution from older pre-Pan African continents (e.g. Dixon and Golombek, 1988; Farahat et al., 2004; Hargrove et al., 2006) and/or plume interaction (e.g. Stein and Hofmann, 1994; Teklay et al., 2002; Farahat, 2006) has been suggested to account for the anomalously high crustal growth rate of the ANS. Recently, Farahat et al. (2007) related such a high growth rate to asthenospheric upwelling due to oblique convergence during precollisional evolution and subsequent mantle delamination during the late- to post-collisional phase. In the Egyptian Eastern Desert, ophiolites are widely distributed in the central (CED) and southern (SED) parts, while entirely lacking from the northern part. In the CED, they cluster in the northern (NCEDO) and southern (SCEDO) segments (Fig. 1). These ophiolites are dismembered bodies composed of mostly serpentinized ultramafics, metagabbros and massive to pillowed metavolcanic rocks. Sheeted dykes are rare in the CED ophiolites and are only found in the SCEDO (e.g. El Sharkawy and El Bayoumi, 1979; Abu El Ela, 1985). Many CED ophiolites are essentially mélanges likely produced by tectonic disruption related to deformation by the post-accretionary Najd fault system, rather than representing a suture zone (Wallbrecher et al., 1993; Abdelsalam and Stern, 1996).

Although a great deal of work has been carried out on the ANS ophiolites over the last 30 years, questions remain as to their origin and tectonic significance. Almost all these ophiolites show clear supra-subduction zone (SSZ) geochemical signatures. Nonetheless, controversy exists on whether they are formed in back-arc (e.g., Amstutz et al., 1984; Khudeir and Asran, 1992; El Sayed et al., 1999; Farahat et al., 2004; Abd El-Rahman et al., 2009a) or forearc (e.g., Stern et al., 2004; Azer and Stern, 2007; Abd El-Rahman et al., 2009b) settings. Most ophiolites may include both MOR- and island arc-types due to the multistage histories of the ophiolite (e.g., Dick and Bullen, 1984; Meffire et al., 1996; Portnyagin et al., 1997; Pearce et al., 2000; Flower and Dilek, 2003; Saccani and Photiades, 2004). This is in agreement with the detailed investigation of the WZO mantle section (Farahat et al., 2010a) which reveals MORB, island arc and boninitic affinities. This observation highlights the need for a comprehensive study of all parts of the Egyptian ophiolites. Recent studies on Egyptian ophiolites (Farahat, 2001; Abd ElRahman et al., 2009a, b; Farahat et al., 2010a,b) pointed out the occurrence of two separate ophiolite types in the CED, both displaying supra-subduction zone (SSZ) geochemical characteristics (Fig. 1). Those clustering in the northern part (NCEDO) show MOR- to island arc tholeiitic-type (IAT) affinities with significant, although volumetrically subordinate, occurrence of boninitic rocks. Ophiolite sequences

Fig. 1. Distribution of ophiolites and ophiolitic mélanges in the central (CED) and southern (SED) parts of the Eastern Desert of Egypt (modified from Shackleton, 1994). The major shear zone (heavy dashed line) separating the CED and SED is from Stern and Hedge (1985). HFSE data (TiO2, Zr, and Y) on volcanic sequences of ophiolites clustering in the northern and southern segments of the CED are from this study, Abdel-Karim et al. (2008), Farahat et al. (2004), and Farahat (2001). The data of the (~ 750 Ma) Wadi El Dabbah island-arc metavolcanics are from Ali et al. (2009). Insert, Neoproterozoic outcrops (white) in the Eastern Desert of Egypt, Sinai, Saudi Arabia (SA), Sudan and Libya. Locations of Fig. 2a and b are indicated.

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in the southern part display more MORB/within-plate geochemical signatures (Farahat et al., 2004). The purpose of this paper is to present new mineralogical and geochemical data on the volcanic sections of the Wizer (WZO) and Abu Meriewa (AMO) ophiolites as representatives of the NCEDO and SCEDO, respectively, in order to discuss their petrogenetic and geotectonic evolution. The results obtained, in conjunction with published geochemical data on CED ophiolites, allow me to propose a new tectono-magmatic evolution model for CED ophiolite formation. This paper is complementary to earlier papers focusing on the mineral composition of WZO and AMO mantle sections (Farahat et al., 2010a; Abu El Ela and Farahat, 2010, respectively). 2. Geologic setting 2.1. Regional geology The Eastern Desert of Egypt is divided into three domains that are structurally and lithologically distinct (Stern and Hedge, 1985; Fig. 1). The northern domain (NED) is characterized by abundant granitoids, minor island-arc metavolcanics, and no ophiolitic rocks. The southern domain (SED) contains diverse lithologies, including gneisses, granitoids, island-arc volcanosedimentary associations and ophiolites that occur as elongated belts (Fig. 1). Of particular interest in the central part of the Eastern Desert (CED), is the occurrence of a complete record of the ANS lithologies. The latter comprise low to medium grade metamorphosed island-arc volcanosedimentary assemblages, dismembered ophiolite sequences that cluster in the northern and southern segments, metagabbros-diorite complexes, and calc-alkaline granitoids formed during the pre-collisional (i.e. island-arc/marginal basin) stage between 750–650 Ma ( Ali et al., 2009; Andresen et al., 2009). Slightly deformed calc-alkaline granodiorites formed during the terrane accretion stage (Stern and Hedge, 1985). On the other hand, a considerable volume of largely undeformed latecollisional calc-alkaline granitoids were emplaced, following crustal thickening associated with terminal oblique collision between fragments of East and West Gondwana and preceding the onset of the crustal extensional phase (~630–590 Ma). A major period of alkaline A-type granites (610–560) and mafic to felsic, high-K, Dokhan volcanics (610–580 Ma) is associated with escape tectonics, transpression, and crustal extension in the post-collisional stage (e.g. Stern and Gottfried, 1986; Stern, 1994; Beyth et al., 1994). Both the WZO and AMO ophiolites lie in the CED and have been studied by earlier workers in the light of the geosynclinal concept (e.g. El Ramly and Al Far, 1955; Sabet, 1961; Noweir et al., 1983). More recently, a marginal basin tectonic setting has been advocated for these ophiolites based on the geochemical characteristics of their crustal sections (e.g. Habib et al., 1987; Khudeir and Asran, 1992). 2.2. Field relations 2.2.1. Wizer area The Wizer area is occupied mainly by a dismembered ophiolite suite folded into a major NW-trending syncline (Khudeir and Asran, 1992). The ophiolitic rocks are represented by abundant ophiolite mélange and a remnant ophiolite nappe (Fig. 2a) which occupies the trough of the syncline and tectonically overlies the ophiolitic mélange. The mélange matrix in the area consists chiefly of regionally metamorphosed layered volcaniclastic turbidites (metaconglomerates, metagreywackes, metamudstones, and schists) enclosing variably sized fragments of ophiolitic rocks dominated by metavolcanic blocks together with some serpentinites and metagabbros (Akaad and Abu El Ela, 2002). The remnant ophiolite nappe comprises serpentinites, metagabbros and metabasalts. Serpentinites constitute the major ophiolite component preserved in the Wizer area. They form N- to NW-trending

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thrust sheets overlying the ophiolitic mélange and form NW-trending narrow belts between metavolcanics and metagabbros of the ophiolite nappe. Serpentinites occur also as small fragments embedded in the mélange matrix. Serpentinites are homogenous and consist chiefly of serpentinized harzburgites. In addition, minor brown to brownish black dunite forms rounded to lensoidal bodies (~ 10–40 cm in size) and small bands parallel to the tectonic fabric; such bodies are occasionally recognized in the lower parts of the serpentinite sheets. In places, dunite bodies are heterogeneous in composition and deformed, containing numerous carbonate veinlets. Adjacent to faults and fractures, serpentinites are highly sheared and intensely altered to talc-carbonates (listwaenite). Metavolcanics and metagabbros occur as fault-controlled small elongated ridges (up to 4 km long) within serpentinites in the trough of the syncline. The ophiolitic metavolcanics and metagabbros are tectonically juxtaposed against the serpentinized mantle section and the Moho is nowhere preserved, i.e. the original internal architecture of the Wizer ophiolite was tectonically disrupted during its emplacement. Metavolcanics are generally massive greyish-green metabasalts. Two dark green boninitic dykes, ~ 0.5 m wide, cross-cut the metabasalt blocks, indicating late-stage development of boninitic magmatism in the area. For this study, the massive metavolcanics of the ophiolite nappe, together with the two dykes and metavolcanic blocks embedded in the mélange, were sampled. 2.2.2. Abu Meriewa area The Neoproterozoic rocks exposed in the Abu Meriewa-Hagar Dungash District (Fig. 2b) comprise an extensive ophiolitic mélange which forms an arcuate belt (~483 km2) trending WNW–ESE and ENE–WSW, together with subordinate calc-alkaline island-arc metavolcanics and an unmetamorphosed mafic layered intrusion (Abu El Ela, 1985; El Amawy, 1997; Abu El Ela and Farahat, 2010). The ophiolitic mélange consists predominantly of imbricate slices of weakly metamorphosed flyschoid metasediments (distal turbidites and olistostromal deposits) and dismembered ophiolitic components of serpentinites and talc-carbonates, metagabbros, sheeted metadiabase dykes and pillow metabasalts (Fig. 2b). The ultramafic rocks of the Abu Meriewa-Hagar Dungash District are represented by serpentinites (of harzburgitic and dunitic parentages) containing chromitite lenses (Abu El Ela and Farahat, 2010). The serpentinites are commonly altered to talc-carbonates along faults and shear zones. They occur as three major thrust sheets (i.e. Abu Meriewa, Hagar Dungash, and Um Huqab sheets) and small fragments embedded in the mélange. The serpentinite sheets are traversed by a number of NW–SE faults that are cut by a major ENE– WSW fault (Fig. 1b). The crustal plutonic assemblage of the AMO is less abundant and occurs as variably sized blocks in the mélange along the eastern, western and central part of the mapped area. It is differentiated into metapyroxenite, layered metagabbros, massive metagabbros, and fine-grained metagabbros. Abu El Ela (1985) observed sheeted metadiabase dyke blocks of varied sized (0.6 to 5 m thick and up to ~1.5 km long) embedded in the metasediments or talc-carbonates along the Wadis Dunqash and Abu Mireiwa. The pillow metabasalts occur as two elongated belts: one along Wadi Um Higlig (~8 km long) and the other one north of Gabal Hager Dunqash (~2 km long). The pillow basalts of both belts are highly sheared and tectonically enclosed within and concordant with the main foliation of the surrounding matrix. The pillows, although less distinct due to extensive deformation, are ovoidal, elongated and lensoidal in shape. 3. Petrography The WZO massive metavolcanics of the nappe are found to be the least altered among the investigated rocks as they exclusively contain

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Fig. 2. Geologic maps of Wizer (a) and Abu Meriewa (b) areas, modified after Khudeir and Asran (1992) and Abu El Ela (1985), respectively.

primary clinopyroxene relics. They are aphyric to porphyritic rocks containing clinopyroxenes partially to entirely altered to actinolite and chlorite (Fig. 3a). Plagioclase forms prismatic albite (An3–8) crystals highly altered to fine-grained epidote (Fig. 3a) and thus are cloudy under the microscope (Fig. 3b). In addition, untwined albite (An3–9) occupies the groundmass. Accessory phases are represented mainly by fine-grained ilmenite partially altered to titanite. WZO metavolcanic blocks in the mélange are entirely altered; they comprise a hornblende + actinolite + albite + epidote + chlorite + titanite + ilmenite mineral assemblage. The two amphiboles (i.e. hornblende and actinolite) form porphyroblasts that are a patchwork of the two phases, characteristic of rocks metamorphosed in greenschist–amphibolite transitional facies. On the other hand, the boninitic dykes are characterized by abundant ferromagnesium

minerals (up to ~80 vol.%) including amphibole (hornblende and actinolite), epidote and chlorite porphyroblasts in a fine-grained groundmass of untwined albite (Fig. 3c). Chalcopyrite and Cr-spinels, partially altered to Cr-magnetite toward rims, are recognized in one of the dykes samples (Fig. 3d). Compared to WZO massive metavolcanics, AMO metavolcanics are more metamorphosed. They contain amphibole + plagioclase + epidote + chlorite ± biotite + titanite. As in the case of WZO metavolcanic blocks in the mélange, amphibole porphyroblasts are subhedral to anhedral pale green actinolite to dark green hornblende. Plagioclase commonly forms euhedral porphyroblasts with albite twinning sometimes destroyed by alteration. Plagioclase composition varies from albite to oligoclase (An4–28). Plagioclase is occasionally enclosed in large amphiboles, possibly the metamorphic product of an

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Fig. 3. Back-scattered electron (BSE) images (a and d) and photomicrographs (b, c, e, and f) of the investigated metavolcanics. a) Euhedral clinopyroxene crystal altered to actinolite (Act) and chlorite (Chl) with clinopyroxene relict (Cpx). Prismatic plagioclase crystals altered to fine-grained epidote (Ep). Untwined albite (Ab) occupies the groundmass. WZO massive metavolcanics. b) Photomicrograph showing cloudy albite crystals (Ab), due to alteration to fine-grained epidote, invading clinopyroxene crystals that are partially to entirely altered to actinolite and chlorite. WZO massive metavolcanics (crossed polarizers). c) WZO boninitic rock consists of abundant amphibole, epidote and chlorite in a finegrained untwined albite groundmass. Crossed polarizers. d) Cr-spinel crystal highly altered to Cr-magnetite (grey) with the preservation of an unaltered core (black). WZO boninites. e) AMO metavolcanic rocks showing twinned to untwined plagioclase (Plg) enclosed in large amphiboles (Amph), possibly the metamorphic product of an original ophitic texture. Crossed polarizer. f) Radiating fibres of plagioclase forming variolitic texture that are encountered near pillow margins. AMO metavolcanic rocks (crossed polarizers).

original ophitic texture (Fig. 3e). Variolitic textures are recognized near pillow margins where the fine, radiating fibres, of plagioclase are enclosed in a glassy groundmass (Fig. 3f). The textural and mineral composition characteristics of the AMO metavolcanics (i.e. actinolite– hornblende coexistence, the albite to oligoclase compositions of plagioclase and the scarcity of chlorite) suggest metamorphism up to greenschist–amphibolite transitional facies (e.g. Farahat, 2008). 4. Analytical techniques Mineral analyses and backscattered electron (BSE) images, in addition to whole-rock XRF analyses, were performed at the Institute of Earth Sciences (Mineralogy and Petrology), Karl Franzens University, Graz, Austria. A JEOL-JSM 6310 scanning electron microscope with Oxford ISIS EDX and a Microspec WDX (15 kv, 6 nA, Na with TAP,

all other elements on EDX) were used for microprobe analyses and BSE imaging. A range of natural and synthetic minerals was used as standards. The analyses were normally done at a magnification of spot mode. The matrix correction was calculated with the ZAF-correction program. Fe2+–Fe3+ redistribution from electron microprobe analyses is made using the charge balance equation of Droop (1987) for estimating Fe3+. Whole-rock major and trace element contents were obtained using a Bruker Pioneer S4 X-ray fluorescence (XRF) spectrometer. Fused glass discs using Li2B4O7 flux were used for major oxides, whereas pressed powder pellets were used for trace element determinations. Loss on ignition (LOI) was determined by heating the powdered samples for 1 h at 1000 °C. The analytical precision and accuracy is better than 3% for major elements and 10% for trace elements.

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Table 1 Whole-rock major (wt.%) and trace element (ppm) analyses of WZO and AMO metavolcanics. Sample

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total Mg# Sc V Cr Co Ni Cu Zn Ga Rb Sr Y Zr Nb Ba Hf La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Th U Sample

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total Mg# Sc V Cr Co Ni Cu Zn Ga Rb Sr Y Zr Nb

WZO massive metavolcanics

WZO blocks in mélange

WZ27

WZ38

WZ41

WZ31

WZ37

WZ33

WZ28

WZ30

WZ31

WZ43

WZ23

WZ45

WZ41

WZ2

WZ14

WZ9

WZ6

46.10 1.81 15.66 11.66 0.17 8.17 9.63 2.74 0.12 0.14 2.91 99.11 60.18 46.9 247 290 83 120 50 81 18 n.d. 320 33.0 129 3.0 33.0 3.0 5.30 15.90 2.69 13.00 3.90 1.62 4.80 0.90 5.40 1.10 3.40 0.49 2.90 0.41 0.1 0.1

47.66 1.64 16.09 10.36 0.16 8.09 9.86 3.32 0.09 0.16 2.00 99.43 62.73 34.6 236 491 57 205 65 74 15.6 n.d. 364 29.6 102 4.0 53.2 2.4

47.62 1.76 15.74 11.01 0.17 8.97 5.86 3.82 0.18 0.16 3.92 99.21 63.73 50.8 295 300 72 120 90 85 17 2.2 167 32.0 146 3.0 42.0 3.4 6.60 15.20 2.62 13.30 4.00 1.49 4.90 0.90 5.50 1.20 3.50 0.51 3.10 0.45 0.1 0.1

48.44 1.41 16.08 10.26 0.17 8.17 9.65 3.43 0.05 0.13 2.32 100.10 63.18 34.9 226 460 60 98 48 78 14.5 n.d. 226 28.1 92 2.0 51.8 2.6

45.42 1.77 16.70 10.75 0.16 8.40 9.29 3.05 0.16 0.17 3.09 98.97 62.77 39.7 238 304 61 149 60 80 15.3 2.4 321 31.0 134 3.0 68.7 1.7

46.51 1.75 15.95 10.75 0.17 8.10 10.38 2.94 0.09 0.17 2.69 99.49 61.89 45.1 243 335 70 140 57 73 16.5 n.d. 428 28.0 114 5.0 55.8 2.7

47.56 1.50 15.38 10.47 0.17 8.38 9.22 2.92 0.12 0.16 2.96 98.84 63.32 38.2 249 456 63 129 60 80 16.9 0.8 375 31.9 112 3.7 23.2 2.9

48.06 1.39 16.25 10.14 0.16 8.21 9.91 3.21 0.09 0.13 2.53 100.07 63.59 40.8 241 368 56 139 69 75 15.6 0.8 235 29.6 106 2.8 12.8 2.5

48.12 1.47 15.96 10.65 0.16 8.69 9.17 3.26 0.06 0.11 2.16 99.81 63.76 38.2 230 471 95 101 38 87 16.4 n.d. 242 30.6 102 2.0 25.6 2.5

47.54 1.83 18.89 10.29 0.16 9.85 4.44 3.05 0.12 0.13 3.48 99.78 67.37 51 305 351 66 123 58 84 17.3 2.4 150 36.0 147 3.2 25.0 3.4

46.76 1.66 15.98 10.64 0.18 8.43 9.54 3.19 0.11 0.22 2.89 99.60 63.08 41.1 258 614 70 168 58 81 16 n.d. 438 32.8 120 4.6 24.3 2.9

46.36 1.73 15.77 10.22 0.16 8.54 10.66 2.85 0.07 0.16 3.23 99.72 64.30 39.3 239 345 78 109 105 81 15.5 n.d. 333 33.7 135 2.9 16.5 3.2

46.79 1.77 16.43 10.55 0.16 8.69 6.82 3.93 0.17 0.18 4.24 99.72 63.98 36.6 234 395 59 159 50 73 17.9 4.1 214 32.6 119 3.0 71.2 2.1

51.23 0.80 17.82 10.74 0.15 4.48 8.02 2.68 0.62 0.32 2.23 99.08 47.05 27.5 437 295 100 96 211 83 25 4.4 421 16.3 97 3.0 318.1 2.8

49.45 0.74 15.21 11.23 0.18 6.75 10.15 3.83 0.16 0.08 1.39 99.17 56.42 45.7 324 83 124 54 34 83 12.6 3 277 16.2 54 2.0 97.0 1.0 2.70 6.60 1.01 5.10 1.60 0.71 2.20 0.40 2.80 0.60 2.00 0.31 2.00 0.28 0.2 0.2

50.40 0.73 15.40 11.87 0.18 5.98 7.58 3.54 0.52 0.14 2.92 99.26 52.07 40.8 350 342 60 117 117 84 14.2 7.2 322 18.4 50 2.2 126.1 1.5 2.62 7.33 1.12 5.66 1.78 0.76 2.44 0.44 3.11 0.67 2.10 0.31 2.22 0.31 0.2 0.1

51.95 0.73 16.23 11.29 0.15 5.18 7.74 2.61 0.54 0.26 2.83 99.51 50.02 32.1 409 354 54 107 194 83 20.3 4.7 405 15.1 93 2.0 342.5 2.6

WZO boninites

AMO LREE-enriched metavolcanics

WZ36

WZ34

AM17

AM19

AM29

AM27

AM33

AM21

AM10

AMO LREE-depleted metavolcanics AM12

AM30

AM22

AM20

AM25

AM11

52.93 0.22 15.66 9.36 0.15 6.08 12.27 1.61 0.06 0.04 1.10 99.48 58.33 43.8 233 133 58 61 93 74 13.5 1 120 4.9 14 1.0

53.30 0.18 14.65 9.75 0.15 7.02 11.76 1.31 0.07 0.04 1.62 99.85 60.83 48.1 240 146 54 68 60 60 16 0.7 114 7.1 17 1.0

50.33 1.36 16.49 8.28 0.12 5.80 6.93 3.16 0.75 0.20 6.47 99.87 60.18 23.7 153 131 36 51 26 71 17.2 15.1 461 22.1 130 7.0

46.17 1.41 15.79 9.02 0.13 11.35 9.29 1.79 0.84 0.28 3.45 99.52 73.07 24.9 163 688 53 283 42 73 14.7 11.7 679 15.3 123 5.0

46.91 1.32 16.02 10.23 0.14 11.33 8.43 1.41 0.86 0.21 2.95 99.81 70.49 23.2 166 732 57 301 44 83 14.9 15.7 717 16.7 117 5.3

49.59 1.11 15.35 6.97 0.14 4.60 9.52 4.18 0.83 0.14 7.45 99.87 58.73 37.1 219 178 33 53 25 71 15.5 11.8 374 25.4 99 5.5

52.70 0.97 14.51 8.26 0.14 8.70 8.81 2.25 0.54 0.17 2.69 99.74 69.43 32.1 197 491 37 73 98 78 16.4 9.5 418 19.3 105 1.0

46.79 1.60 16.68 9.33 0.13 9.65 9.28 2.51 0.46 0.31 2.75 99.49 69.05 23.8 178 496 48 206 20 81 16.1 7.5 588 17.3 121 5.0

49.05 1.60 14.17 12.81 0.18 6.24 9.41 3.18 0.10 0.13 2.87 99.73 51.22 60.5 383 209 49 54 28 102 17.3 0.4 102 41.6 89 2.0

50.35 1.75 13.65 13.50 0.18 7.32 6.53 3.55 0.06 0.15 2.49 99.52 53.89 62.9 422 233 46 65 28 106 16.7 0.9 116 44.6 107 2.0

48.85 1.44 13.67 12.17 0.18 6.94 10.01 2.88 0.13 0.12 3.02 99.41 55.14 51.5 467 220 50 52 26 116 17.4 0.2 109 50.5 105 3.0

49.58 1.52 13.14 14.75 0.17 7.90 7.24 2.64 0.08 0.12 2.61 99.75 53.60 58.4 521 260 45 63 35 120 17 0.3 123 50.2 107 2.0

48.37 1.15 13.90 12.08 0.19 8.74 8.92 2.22 0.26 0.08 3.55 99.46 60.94 47.1 308 333 50 90 47 98 15.3 5 239 33.5 64 1.0

49.15 1.12 13.77 11.99 0.18 8.06 8.26 3.01 0.21 0.10 3.87 99.71 59.19 52.7 315 299 46 91 42 86 13.2 3.5 247 28.9 47 1.0

50.58 2.08 12.84 13.17 0.23 8.50 7.50 2.90 0.09 0.17 1.65 99.71 58.20 62.2 472 94 53 59 5 117 15.4 0.5 61 52.0 117 1.0

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Table 1 (continued) Sample

Ba Hf La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Th U

WZO boninites

AMO LREE-enriched metavolcanics

WZ36

AM17

AM19

AM29

AM27

AM33

AM21

AM10

AM12

AM30

AM22

126.3 3.8 15.80 38.70 5.20 22.20 5.30 1.70 4.93 0.77 4.93 1.02 2.89 0.40 2.30 0.31 0.7 0.3

189.1 0.0

172.2 3.0

617.6 3.0

357.3 0.7

132.6 1.4 11.10 29.10 4.38 17.70 4.30 1.51 4.10 0.60 3.60 0.70 2.10 0.28 1.70 0.25 0.5 0.2

40.7 2.9

85.9 2.3

21.8 2.5

53.3 2.5

28.0 0.4 1.31 3.03 0.39 1.65 0.44 0.16 0.48 0.09 0.74 0.17 0.62 0.11 0.78 0.13 0.3 0.4

WZ34 22.0 0.5 1.50 3.60 0.45 1.90 0.50 0.22 0.70 0.10 1.00 0.20 0.80 0.14 0.90 0.15 0.3 0.2

AMO LREE-depleted metavolcanics AM20 72.4 1.7 2.20 6.70 1.22 7.30 2.60 0.93 3.50 0.70 4.70 1.00 3.30 0.50 3.10 0.45 0.1 0.1

AM25 77.9 0.9

AM11 28.7 5.0 3.10 11.30 2.18 13.20 4.70 1.67 6.30 1.20 8.30 1.80 5.90 0.88 5.40 0.82 0.1 0.2

Major and trace element determined by XRF; selected REE, Th and U analyzed by ICP-MS; Mg# = molar Mg/(Mg + Fe) ⁎ 100; Fe2O3 = 0.2FeO; n.d. = not detected.

Selected samples were further analyzed for trace elements including Rare Earth Elements (REE), and Th that were not determined by XRF, using inductively coupled plasma-mass spectrometry (ICPMS) at the Activation Laboratories, Ontario, Canada. Detailed analytical methodology can be found in the ACTLABS Group website (http:// www.actlabs.com). It is important to mention that a close match is observed between trace elements determined by both XRF and ICP-MS techniques including those of low values such as Nb and Hf. 5. Geochemistry SiO2 contents (Table 1) of both WZO and AMO metavolcanic rocks confirm that they are dominantly basalts (SiO2 = 45.42–52.7 wt.%); except for the two WZO dyke samples, which display boninitic affinity. According to Le Bas (2000) boninites are volcanic rocks that contain N8 wt.% MgO, N52 wt.% SiO2 and b0.5 wt.% TiO2. The SiO2 (52.93 and 53.3 wt.%) and TiO2 (0.22 and 0.18 wt.%) contents of the two dyke samples match those of boninitic rocks. The lower MgO contents (7.02–6.08 wt.%) of the two dyke samples with respect to typical boninites are, however, most likely linked to high mobility of MgO in hydrothermal fluids (e.g. Bach et al., 2003). Crawford et al. (1989) subdivided boninites into high- and low-Ca types (CaO/ Al2O3 N 0.75 and b0.75, respectively). Accordingly, the investigated boninites can be classified as high-Ca boninites (CaO/Al2O3 = 0.78 and 0.80). The boninitic signature of the two dyke samples is further indicated by their extreme depletion in high field strength elements (HFSE) such as Zr (13.9–17 ppm) and Y (5–7 ppm). In the SiO2 vs. Nb/ Y classification diagram (Fig. 4a) the WZO and AMO metavolcanic rocks plot in the subalkaline basalt field, whereas the two boninitic

samples straddle the subalkaline basalt–andesite boundary. On the Zr vs. Y diagram of Barrett and MacLean (1994), most of the WZO and AMO rocks occupy the tholeiitic field; with a few samples from both ophiolite sequences extending into the transitional field (Fig. 4b). Both WZO and AMO metavolcanics define a low-K suite with average K2O contents that vary from 0.18 to 0.40 wt.%, respectively. The WZO metavolcanics can be divided into high- and low-Ti basalts, in addition to the very low-Ti boninites. The massive metavolcanics have high TiO2 content (1.39–1.83 wt.%) similar to MORB (Pearce, 1982), whereas the metavolcanic blocks in the mélange are characterized by low TiO2 content (0.73–0.8 wt.%) like island-arc tholeiites (IAT). The AMO metavolcanics have TiO2 content between 0.97 and 2.08 wt.%. Major and trace element concentrations of the investigated samples are plotted against Zr, which is not disturbed by modest alteration and metamorphism and which also serves to monitor fractionation of mafic melts (Fig. 5). Both WZO and AMO metavolcanic suites display two chemically distinguishable groups, implying different fractionation trends or degrees of partial melting of the same source and/or different source compositions. The generally good correlations between Zr and both transition metals (e.g. V and Sc) and HFSE (e.g. Y and Ti), contrary to most major element and large ion lithophile elements (LILE) (e.g. Sr), indicate that the HFSE have been slightly mobilized by alteration, as will be elaborated more fully latter. The WZO massive metavolcanics have flat to slightly enriched LREE chondrite-normalized REE patterns [La/YbCN (i.e. chondritenormalized ratio) = 1.31 and 1.53]. Their HREE abundances are comparable to those of MORB, with 16–24 times chondrite abundances (Fig. 6a). The WZO metavolcanic blocks in the mélange have,

Fig. 4. Geochemical classifications of the ophiolitic metavolcanic rocks. a) SiO2 vs. Nb/Y plot (after Winchester and Floyd, 1977). b) Zr vs. Y diagram (after Barrett and MacLean, 1994).

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E.S. Farahat / Lithos 120 (2010) 293–308

2.5

(b)

52

SiO2 (wt.%)

TiO 2 (wt.%)

2

54

(a)

WZO massive metavolcanics WZO blocks in melange WZO boninites AMO LREE-depleted rocks AMO LREE-enriched rocks

1.5 1

50 48 46

0.5

44

0

(d) Al2O3 (wt.%)

MgO (wt.%)

(c) 8

4

18

16

14

12

(e)

(f)

0.8

14

CaO/Al2O3

Fe3O3 (wt.%)

0

12 10

0.6

0.4 8 6 300

200

100

400

200

0

0

(i) 400

(j)

60

Sc (ppm)

V (ppm)

(h)

600

Cr (ppm)

Ni (ppm)

0.2

(g)

200

40

20

0

0

(k)

(l)

Sr (ppm)

Y (ppm)

600 40

400

20 200

0

0

40

80

Zr(ppm)

120

160

0 0

40

80

120

Zr(ppm)

Fig. 5. Variation diagrams for selected major and trace elements vs. Zr for the WZO and AMO metavolcanics.

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301

Fig. 6. Chondrite-normalized REE patterns (a and b) and N-MORB-normalized multielement diagrams (c and d) for the WZO (a and c) and AMO (b and d) metavolcanics. Normalizing values are taken from Sun and McDonough (1989).

however, overall lower total REE abundances compared to the WZO massive metavolcanics. The two dyke samples display markedly U-shaped REE patterns typical of boninitic rocks (e.g. Hickey and Frey, 1982; Cameron et al., 1983; Meffire et al., 1996). The AMO rocks exhibit two very different REE patterns (Fig. 6b). One group of samples displays LREE-depleted (La/GdCN = 0.43–0.55) and flat HREE patterns (Gd/YbCN = 0.93–0.97). The other group shows highly fractionated LREE/HREE patterns (La/YbCN = 4.68– 4.93; La/GdCN = 2.35–2.78; Gd/YbCN = 1.78–1.99). The AMO metavolcanics can thus be categorized into LREE-enriched and LREEdepleted lava types, which correspond to the two trends for the AMO shown in the variation diagrams (Fig. 5). Different symbols are thus used for the two AMO lava types in the various geochemical diagrams, although all of them were not analyzed for REE. The MORB-normalized trace element diagrams of selected samples representing different magma types of WZO and AMO metavolcanics are shown in Fig. 6c and d. Both WZO and AMO samples show high abundances of fluid-mobile LILE such as LREE, K, Ba, Sr, Th and U with respect to fluid-immobile HFSE, coupled with negative Nb-anomalies, pointing to subduction-modified mantle sources (e.g. McCulloch and Gamble, 1991). It is important to note that the two boninitic samples exhibit the highest LILE enrichment among the WZO rocks, implying a strong contribution from subduction components. Systematic behavior of LILE and negative Nb-anomalies in combination with the immobile HFSE is good evidence that the generally overabundance of enrichment in fluid-mobile elements is a primary characteristic of both the WZO and AMO crustal sections. 6. Clinopyroxene and spinel chemistry Both primary clinopyroxene and spinel chemical compositions reflect the composition of the magmas from which they crystallize. It is thus widely accepted that their compositions represent a suitable indicator of the magmatic affinity of basalts from different tectonic settings (e.g. Leterrier et al., 1982; Barnes and Roeder, 2001) and from different ophiolitic types (e.g. Beccaluva et al., 1989, Stern et al., 2004). Consequently, preserved primary clinopyroxenes and spinels from the WZO massive metavolcanics of the nappe and boninitic dykes, respectively, were analyzed and representative chemical data are presented in Table 2. Clinopyroxene compositions are uniform and plot in the augite field of the wollastonite–enstatite–ferrosilite classification

diagram (Fig. 7). They are characterized by high Ti-contents (0.91– 2.06 wt.%, 1.42 wt.% on average) and by moderately high Mg# [100 Mg/(Mg + Fe2+), atomic ratio; 61.1–87.3, mean = 75.6]. Using the discrimination diagram of Beccaluva et al. (1989) clinopyroxenes from the massive metavolcanics reveal a close affinity with clinopyroxenes of normal-MORB, in harmony with their whole-rock chemistry. On the other hand, the Cr# [100Cr/(Cr + Al), atomic ratio] and Mg# of unaltered spinel cores from the boninitic dykes (83–84 and 32–52) are comparable to those from dunite bodies found in the WZO mantle section (81–87 and 36–64; Farahat et al., 2010a) which are residues after boninite melt formation. 7. Discussion 7.1. Alteration process and element mobility Generally, the ophiolitic rocks of the Eastern Desert of Egypt underwent two metamorphic events, resulting in greenschist to greenschist–amphibolite transitional facies metamorphism (e.g. Farahat, 2008): an earlier low-temperature seafloor hydrothermal event and a later higher temperature regional metamorphic event. Petrographic investigations of the WZO and AMO samples revealed that the rocks have undergone variable degrees of alteration. Given the preservation of primary clinopyroxene in the WZO massive metavolcanics, these rocks most likely suffered exclusively seafloor hydrothermal alteration and escaped later regional metamorphism. By contrast, the metavolcanic blocks in the mélange and the AMO highly sheared metavolcanics are entirely altered. Consequently, any redistribution of elements during alteration and metamorphism needs to be evaluated before any further interpretation of the chemical data. Loss on ignition (LOI; anticipated to consist mainly of H2O) has been used to monitor the extent of element redistribution during alteration (Polat et al., 2002; Polat and Hofmann, 2003). With the exception of two AMO samples (AM17 and AM27, Table 1) which have high LOI values (6.47 and 7.45 wt.%) all WZO and AMO rocks have LOI b4 wt.%, suggesting moderate alteration. Generally, an agreement exists that some transition metals (e.g. Ti, Cr, V, Sc), REE, HFSE, as well as Th are relatively immobile during low-temperature alteration (Gélinas et al., 1982; Staudigel et al., 1996; Polat and Hofmann, 2003). This seems correct for the investigated rocks where HFSE (e.g. Nb, Ti, Zr, Hf, Y) and REE from both WZO and AMO display

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Table 2 Representative microprobe analyses of clinopyroxenes (formulae based on 6 oxygen) and Cr-spinels (formulae based on 4 oxygen) from the WZO massive metavolcanics and boninites, respectively. Clinopyroxenes in the massive metavolcanics

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO MgO CaO Na2O ZnO Total Si Ti Al Cr Fe3+ Fe2+ Mn Mg Ca Na Zn Mg# Cr#

Cr-spinels in the boninites

Cpx 3

Cpx16

Cpx24

Cpx27

Cpx30

Cpx32

cpx1

Cpx 7

Sp11

Sp12

51.75 1.03 3.59 0.28 0.44 7.38 0.20 15.66 19.91 0.39 – 100.64

50.05 1.26 2.66 0.00 2.39 10.13 0.43 13.54 18.96 0.43 – 99.85

50.42 1.15 3.48 0.15 3.32 6.05 0.16 16.05 19.37 0.34 – 100.48

49.39 1.58 3.60 0.04 3.49 8.64 0.23 14.19 18.99 0.43

49.22 1.61 3.22 0.03 2.92 10.60 0.38 12.96 18.59 0.50 – 100.02

50.53 1.21 3.60 0.13 1.95 7.88 0.22 15.52 18.83 0.35 – 100.23

49.25 2.08 4.09 0.03 0.00 12.88 0.27 11.33 18.86 0.58 – 99.37

48.93 1.89 5.30 0.05 1.02 7.80 0.20 13.94 19.44 0.53 – 99.10

0.46 0.18 7.76 59.19 3.58 17.06 0.23 10.53 0.17 – 0.17 99.39

0.31 0.07 7.67 58.36 3.25 20.81 0.79 7.41 0.27 – 0.16 99.12

1.898 0.028 0.155 0.008 0.012 0.226 0.006 0.856 0.782 0.028 – 79.11 4.91

1.887 0.036 0.118 0.000 0.068 0.319 0.014 0.761 0.766 0.031 – 70.46 0.00

– 100.58

1.858 0.032 0.152 0.004 0.092 0.186 0.005 0.881 0.765 0.024 – 82.57 2.56

1.842 0.044 0.158 0.001 0.098 0.269 0.007 0.789 0.759 0.031 – 74.57 0.63

1.860 0.046 0.143 0.001 0.083 0.335 0.012 0.730 0.753 0.037 – 68.54 0.69

1.871 0.034 0.157 0.004 0.054 0.244 0.007 0.857 0.747 0.025 – 77.84 2.48

1.879 0.060 0.184 0.001 0.000 0.410 0.009 0.644 0.770 0.043 – 61.10 0.54

1.834 0.053 0.234 0.001 0.029 0.244 0.006 0.779 0.781 0.039 – 76.15 0.43

0.015 0.005 0.306 1.564 0.090 0.477 0.007 0.525 0.006

Sp8 0.43 0.11 7.82 59.66 3.53 17.69 0.11 10.11 0.24 – 0.34 100.14

0.011 0.002 0.310 1.581 0.084 0.597 0.023 0.379 0.010

–

–

0.004 52.40 83.64

0.004 38.83 83.61

0.014 0.003 0.307 1.572 0.088 0.493 0.003 0.502 0.009 – 0.008 50.45 83.66

Sp5 0.43 0.08 7.40 58.69 2.53 21.93 0.71 6.96 0.17 – 0.18 99.07 0.015 0.002 0.301 1.595 0.066 0.632 0.021 0.358 0.006 – 0.004 36.16 84.12

Sp2 0.26 0.17 7.71 58.11 3.08 23.36 1.19 5.65 0.21 – 0.07 100.07 0.009 0.004 0.313 1.581 0.080 0.672 0.035 0.290 0.008 – 0.002 30.15 83.47

Mg# = 100 ⁎ Mg/(Mg + Fe2+); Cr# = 100 ⁎ Cr/(Cr + Al).

coherent MORB- and chondrite-normalized patterns (Fig. 6), indicating that the concentrations of these elements were not modified extensively during post-magmatic alteration. In contrast, alkaline earths such as Mg, Ca, Sr and Ba and alkali metals such as Na, Rb and K are commonly mobilized during alteration (e.g. Winchester and Floyd, 1977; Polat and Hofmann, 2003). The relative variations of these elements with respect to Zr and in the MORB-normalized patterns (Figs. 5 and 6) signify that they may be affected by some mobilization during alteration and therefore particular caution should be applied in their use. However, the rather systematic behavior of LILE and negative Nb-anomalies in combination with the immobile HFSE is evidence that the LILE have been moderately mobilized by the alteration and thus can be used, cautiously, to quantify their enrichment with respect to HFSE.

7.2. Low pressure fractional crystallization As mentioned above, the investigated rocks are almost entirely altered and therefore the fractional crystallization processes cannot be evaluated in detail. However, the generally high Mg# (Table 1) of the WZO massive metavolcanics (60.18–67.37; 66.38 on average) and AMO rocks (51.22–73.07; 61.01 on average), together with their enrichment in Cr (290–614 and 94–731 ppm, respectively) and Ni (98–205 and 50–300 ppm, respectively), indicates fairly primitive compositions (Table 1). Nonetheless, the reasonably good correlation within each of the investigated rock groups between Zr and many major and trace elements (Fig. 5) indicates fractional crystallization was an important process. This is further suggested by the Cr vs. Y diagram (Fig. 8), where all samples show trends of Cr

Fig. 7. Clinopyroxene composition of the WZO massive metavolcanics plotted in (a) the wollastonite (Wo)–enstatite (En)–ferrosilite (Fs) classification diagram of Morimoto (1988) and the (b) TiO2–SiO2/100-Na2O tectonic discrimination diagram of Beccaluva et al. (1989).

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7.3. Mantle source characteristics and partial melting

Fig. 8. Cr vs. Y covariations of WZO and AMO metavolcanics (modified after Pearce, 1983) in relation to mantle source compositions and melting trends for incremental batch melting (after Murton, 1989). The percentages of partial melting are indicated. M1: calculated MORB source; M2: residue after 20% MORB melt extraction; M3: residue after 12% melt extraction from M2.

decreasing with increasing Y. This antivariation is expected from the fractional crystallization of olivine, Cr-spinel and clinopyroxene (Pearce, 1982). The early crystallization of Cr-spinel and olivine is confirmed by the decrease of Ni and Cr with increasing Zr. Yet, the Ni contents of the WZO massive metavolcanics contrary to the WZO metavolcanic blocks in the mélange and AMO, first increase then decrease with increasing Zr (Fig. 5). This most likely indicates early crystallization of Cr-spinel followed by olivine, which is the main host of Ni. Furthermore, the decrease of SiO2, Sr and Al2O3 coupled with increase of MgO, TiO2, V, Y, Sc and CaO/Al2O3 with increasing Zr of the WZO massive metavolcanics and the LREE-depleted type of AMO lava likely indicates the early crystallization of plagioclase relative to clinopyroxene. In contrast, the decrease of Fe2O3, Sc, V, and Y, coupled with enrichment in Sr, of the LREE-enriched type of AMO samples likely signifies the early crystallization of clinopyroxenes with respect to plagioclase (Fig. 5). On the other hand, the positive correlation between Ti and Zr suggests that the crystallization of Fe–Ti oxides was not significant. The limited number of samples further complicates evaluating the role of fractional crystallization in the magmatic evolution of the WZO mélange metavolcanic blocks and the boninitic dykes. However, one of the main characteristics of these rocks is their lesser Cr and Ni contents compared to the WZO massive metavolcanics (Fig. 5). This can be attributed either to a lower abundance of these elements in the mantle source, or to early fractionation of Cr- and Ni-bearing phases (i.e. olivine, Cr-spinel and clinopyroxene). Ni and Cr abundances in mantle sources are not extensively modified during the progressive mantle source depletion (Pearce, 1983) inferred, in the next section, for the generation of the WZO rocks. Consequently, the low Cr and Ni abundances in the WZO metavolcanic blocks in the mélange and the boninitic dykes can be accounted for by the early fractionation of Cr- and Ni-bearing phases. This is further supported by lower Mg# (47.3–56.4 and 58.3–60.8 for blocks in the WZO mélange and boninites, respectively) compared to the WZO massive metavolcanics (Mg# = 60.2–67.4).

From the overall geochemical characteristics, the metavolcanic rocks from both the WZO and AMO represent different magma types that could be generated from various mantle sources and/or from different degrees of partial melting. Nonetheless, all the investigated rocks exhibit MORB-normalized trace element patterns (Fig. 6) characterized by variable enrichment of LILE relative to HFSE, coupled with negative Nb-anomalies, which are generally explained as due to melting of mantle that was metasomatised by fluids released from a subducted slab (Pearce, 1982; Tatsumi and Kogiso, 2003). LILE and/or LREE enrichment are the subduction-derived components, whereas HFSE are inherited from the mantle wedge due to their low solubility in aqueous fluids (Hawkesworth et al., 1993; Pearce and Peate, 1995; John et al., 2004). HFSE have been used, therefore, to depict the composition, and enrichment or depletion history of the mantle wedge (Woodhead et al., 1993). Thus, Zr/Yb and Y/Yb vs. Nb/Yb diagrams (Fig. 9a, b) are used to show how depleted is the mantle source (Pearce and Peate, 1995; Pearce, 2008). Yb is used as the normalizing factor in these ratios to eliminate variations that may result from partial melting and fractional crystallization in the spinel peridotite stability field. Fig. 9a and b show that all the investigated rocks plot within the mantle array, which is defined by mantle-derived oceanic basalts. The WZO rocks cluster around the MORB source, whereas the AMO metavolcanics show some source heterogeneity. The LREE-depleted lava type extends toward lower Nb/Yb values, indicating derivation from a more depleted mantle source. On the other hand, the LREE-enriched AMO lava type lies midway between N-MORB and OIB sources, implying an enriched mantle source. The Zr values from all the rocks fall in the mantle array, signifying the limited role of melting of subducted sediments in the evolution of these rocks (Pearce and Peate, 1995). In contrast to HFSE, LREE (La) and Th are more effective in discriminating depleted and enriched mantle sources metasomatised by subduction-derived fluids (Pearce, 1983, 2008). Again, all the investigated rocks fall within the mantle array on the La/Yb vs. Nb/Yb diagram (Fig. 9c), indicating that the LREE were not extensively added to the mantle wedge from the subducted slab. Similar limited or insignificant contribution of LREE from the subducted slab is observed in many SSZ rocks (e.g. Klein and Karsten, 1995; Yumul, 1996; Beccaluva et al., 2005; Saccani et al., 2008). Note that LREE-enriched AMO metavolcanics plot within the mantle array but towards OIB basalt (Fig. 9c) indicating an enriched mantle source rather than significant LREE enrichment from the subducted slab. In contrast, Th of the WZO metavolcanic sequence shows progressive SSZ metasomatism. Fig. 9d shows that WZO massive metavolcanics plot in the lower boundary of the mantle array near typical MORB values indicating insignificant contribution from the subducted slab to the mantle wedge. On the other hand, the WZO metavolcanic blocks in the mélange lie close to the upper boundary of the mantle array, whereas the boninitic rocks lie above the mantle array in the overlap area between oceanic and continental arcs fields, suggesting modest and strong influence of subduction zone fluids, respectively. As the boninitic magmas originated from ultra-refractory, incompatible trace element depleted sources, the trace element budget, including the characteristic U-shaped REE patterns, is controlled by subduction components (e.g. Hickey and Frey, 1982; Cameron et al., 1983, König et al., 2010). The AMO lava types straddle the upper boundary of the mantle array (Fig. 9d), implying regular and slight SSZ metasomatism as compared to WZO. On plotting a compatible (Cr) vs. an incompatible (Y) element (Fig. 8), it is possible to estimate the degree of partial melting and the composition of mantle sources (Pearce, 1983). In Fig. 8, three possible mantle sources, in accordance with the model of incremental batch melting starting from a single source proposed by Murton (1989), are

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Fig. 9. Zr/Yb, Y/Yb, and La/Yb vs. Nb/Yb (a, b and c, respectively) for the WZO and AMO metavolcanics. The MORB-OIB (mantle) array is from Green (2006). (d) Th/Yb vs. Nb/Yb (after Pearce and Peate, 1995). Normal- and enriched-MORB sources are from Sun and McDonough (1989). MORB and OIB represent mid-oceanic ridge basalt and ocean island basalt mantle sources, respectively.

considered for the genesis of the WZO and AMO crustal rocks. The WZO massive metavolcanic rocks follow the N-MORB fractionation trend that intersects the MORB source (M1) calculated according to Pearce (1983) at about 20% partial melting. However, assuming that the WZO metavolcanic mélange blocks originated from a MORB mantle source similar to that of the massive metavolcanics, the possible trend of fractional crystallization intersects the melting path at an implausibly high degree of partial melting (≥40%). Alternatively, they may represent ~ 15% partial melting of a mantle source that previously experienced 20% MORB melt extraction (M2). The subparallel REE patterns of the WZO massive metavolcanics and metavolcanic blocks in the mélange, together with the lower absolute REE concentration of the latter, further support this interpretation. Similarly, assuming that the WZO boninitic dykes originated from a source that had previously experienced only MORB extraction (i.e., source M2), this would require an extremely high degree of partial melting (N40%). As an alternative, these rocks may derive, more plausibly, from ~30% partial melting of the relatively more depleted harzburgitic source (M3) calculated by Murton (1989) as the residue after about 12% melt extraction from M2. Such high melting conditions are in accordance with the model proposed by Falloon and Danyushevsky (2000) that high-Ca boninites represent very high temperature melts (1400–1500 °C) fluxed with 1–2% H2O. The requirement of very high thermal conditions has been suggested by many authors (e.g. Duncan and Green, 1980; Beccaluva and Serri, 1988) to explain the genesis of similarly depleted melts through a multiple-stage melting model. The WZO boninitic dykes show lower HREE and variable LREE enrichment compared to the WZO metavolcanic rocks, which could be related to the combined effect of partial melting of a progressively more refractory mantle source and their variable enrichment in the most incompatible elements (LILE and LREE) from subduction-derived hydrous fluids (Beccaluva and Serri, 1988). Such unique petrogenetic conditions, i.e. ultra-depleted mantle source, abundant water, and an abnormally high geothermal gradient, are most commonly found in the forearc regions of intra-oceanic convergent margins when subduction begins (Crawford et al., 1989). From the above discussion, it is clear that the origin of the different lava types in the WZO volcanic sequence is in agreement with a

multiple-stage melting model in which lower massive metavolcanics of more MORB-like affinities were generated by 20% partial melting of an undepleted lherzolitic (MORB) mantle source leaving, as a residue, a slightly depleted harzburgitic mantle composition. This, in turn, can be considered as the likely mantle source of the metavolcanic blocks in the mélange (for ca. 15% partial melting). The consequent melt extraction left a highly depleted harzburgitic residue, which can be the source for the boninites corresponding to 30% partial melting. The final boninitic-type melt extraction left a residue which is dunitic composition. This petrologic scenario for the origin of the WZO from a mantle source progressively depleted by partial melting is in harmony with that suggested by Farahat et al. (2010a) on the basis of the chemical composition of Cr-spinels in the ultramafic rocks of the WZO mantle section. This is composed of slightly depleted harzburgites and highly depleted harzburgite containing small dunite bodies, which are the residues after MORB, IAT, and boninite melts formation, respectively. The close similarity between unaltered core compositions of Cr-spinels from both the boninitic dykes (Cr# = 83–84) and dunitic bodies (Cr# = 81–87; Farahat et al., 2010a), along with the subordinate distribution of boninitic rocks, provides direct evidence for the validity of the proposed model. On the other hand, the MORB-normalized patterns of the WZO metavolcanic rocks indicate that their mantle sources underwent LILE enrichment by subduction zone-derived fluids. These enrichments likely vary from weak to modest to strong (Fig. 6) for the massive metavolcanics and metavolcanic blocks in the mélange and the boninitic dykes, respectively. As mentioned previously, the different variation trends and REE patterns for the AMO metavolcanics (Figs. 5 and 6) are consistent with mantle source heterogeneity, i.e. depleted to slightly enriched sources. However, on the Cr vs. Y diagram (Fig. 8) the line of fractionation of the AMO LREE-depleted lava types, intersects the partial melting path of the MORB source at about 15%. Moreover, the line of fractionation of the LREE-enriched lava type almost follows those of the WZO metavolcanic blocks in the mélange, implying about 15% partial melting of a depleted source that had previously experienced MORB extraction. Yet, the Nb/Yb ratios for these samples (Fig. 9) argue for an enriched, rather than depleted, mantle source.

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7.4. Tectonic settings and regional tectono-magmatic implications The data presented above indicate that each of the WZO and AMO crustal sections show geochemical characteristics which can be referred to apparently contrasting tectonic settings. Using the standard TiO2 vs. Zr and V vs. Ti tectonic discrimination diagrams (Fig. 10), the WZO massive metavolcanics plot in the overlap areas between MORB and both the within-plate and volcanic arc fields (Fig. 10a) or display Ti/V ratio between 35 and 45 typical of MORB and/or back-arc basin basalts (Fig. 10b). In contrast, the WZO metavolcanic blocks in the mélange lie in the IAT fields in both diagrams. Owing to their depletion in HFSE, the two boninitic samples lie outside the designated fields of Fig. 10a, and display very low Ti/V ratios (b10) characteristic of boninitic rocks (Fig. 10b). Such transition MORB-IAT-boninitic magmatic affinities are considered characteristic of ophiolites formed in a protoarc (forearc) tectonic environment (Casey and Dewey, 1984; Crawford et al., 1989; Shervais, 2001; Dilek et al., 2008) similar to the forearc crust of the Izu–Bonin–Mariana (IBM) arc system (Stern and Bloomer, 1992). These geochemical characteristics are generally shared by the NCEDO volcanic sections (Figs. 1 and 10), where significant boninitic rocks, although volumetrically subordinate, are exclusively observed (e.g. Farahat, 2001; El Sayed et al., 1999). In contrast to what has been observed in the WZO and generally in the NCEDO, rocks with boninitic affinity are entirely lacking in the SCEDO (Farahat et al., 2004; Abd El-Rahman et al., 2009a). The AMO metavolcanic rocks fall mainly in the overlap areas between MORB and both the within-plate and island-arc fields (Fig. 10a). Their MORB to island-arc characteristics are further evidenced by the Ti vs. V discrimination diagram (Fig. 10b). Such MORB/within-plate to island-arc geochemical characteristics, together with the absence of boninitic rocks, the inferred source heterogeneity, and the extensional tectonic regime manifested by the sheeted dykes (e.g. Abu El Ela, 1985; El Sharkawy and El Bayoumi, 1979), are characteristic of ophiolites formed in back-arc basins (e.g. Pearce et al., 1984; Gribble et al., 1998; Sinton et al., 2003). The compositional variations of backarc basin lavas have been accounted for by source heterogeneity related to variable mixing of fertile mantle, melt-depleted mantle and slab-derived components (Martinez and Taylor, 2003; Kuzmichev et al., 2005), which is the case for the AMO and the SCEDO in general (Figs. 1 and 10). This transition in geotectonic affinities of the CED ophiolites from forearc oceanic crust in the north to back-arc oceanic crust in the south, implies that this part of the Eastern Desert could represent an arc–back-arc system, likely in analogy to the Marina arc system, that was formed above a southeast-dipping subduction zone (in modern coordinates). This is further confirmed by the dominance of mature mafic to felsic island-arc metavolcanics, the absence of a crustal break

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and major shear zones in the CED, the overall enrichment in the HFSE of ophiolites from north to south (Figs. 1 and 10), and the separation of the CED from the SED by a major shear zone (Fig. 1). Geochronological data for the ophiolites from the CED, although few, yield rather conflicting results. The Wadi Ghadir ophiolite (Fig. 1) in the extreme southern part of the CED (a zircon 207Pb/206Pb age of 746 ± 19 Ma; Kröner et al., 1992) is younger than the Meatiq dome cover (a zircon 207Pb/206Pb age of 788 ± 13 Ma; Loizenbauer et al., 2001) in the northern part (Fig. 1). Recently, Andresen et al. (2009) presented, however, a U–Pb TIMS zircon age of 736.5 ± 1.2 Ma for a gabbro sample from the Fawakhir ophiolite, located a few kilometres SW of the Meatiq dome. These ages generally overlap the ANS islandarc stage (~ 770–720 Ma; Stern and Hedge, 1985). Ali et al. (2009) reported a U–Pb SHRIMP age of ~ 750 Ma for the CED Wadi El Dabbah island-arc metavolcanics (Fig. 1). Such similar ages between islandarc metavolcanics and ophiolites from the CED are consistent with the proposed model for the CED crust evolution. Spatial variations in incompatible element concentrations (e.g., Ti, Y, and Zr among others, Figs. 1 and 10) across the forearc, through arc to back-arc observed in the CED occur in present day arc–back-arc system (e.g., Taylor et al., 1992; Hochstaedter et al., 2001) and also in the Eritrean Neoproterozoic basement, southern ANS (Teklay, 2006). Such across-arc variations in concentrations of incompatible elements have been accounted for by either different degrees of partial melting of a homogenous source or variations in source region chemistry (Woodhead et al., 1993; Tatsumi and Eggins, 1995; Hochstaedter et al., 2001). The above-mentioned petrogenetic interpretations are in accordance with the second interpretation, i.e. differences in source chemistry. A similar vertical and lateral transition from more MORB to IAT magmatism has been described from the Middle Jurassic Tethyan ophiolites in the Balkan Peninsula (e.g. Dilek and Flower, 2003; Beccaluva et al., 2005; Saccani et al., 2008; Dilek et al., 2008). These authors related these transitional affinities to generation in subduction rollback cycles during the closing stages of basins prior to terminal continental collision. However, the Neo-Tethyan ophiolites are not associated with an evolved magmatic arc, implying formation in a narrow ocean basin, where subduction to N ~100 km (needed to form a magmatic arc) did not occur prior to ophiolite emplacement (R. J. Stern, pers. commun.). Consequently, the recognition of fully developed island-arc metavolcanic sequences having tholeiitic to calc-alkaline signatures, pyroclastic rocks and felsic differentiates, together with ophiolites distribution over ~250 km (Fig. 1), in the CED argues against the subduction rollback model. As an alternative, the subduction zone infancy model of Stern and Bloomer (1992) and Stern (2004), similar to the Eocene Marina arc system, can explain the above documented geologic features of ophiolite sequences occurring in the CED. According to this tectonic

Fig. 10. TiO2 vs. Zr (a) and V vs. Ti/100 (b) tectonic discrimination diagrams of Pearce (1982) and Shervais (1982), respectively. Compositional fields for NCEDO (Fawkheir and Muweillih ophiolites) and SCEDO (Beririq, Gabal Ghadir and Wadi Ghadir ophiolites) are from Farahat (2001) and Farahat et al. (2004). BABB = back-arc basin basalt; IAT = islandarc tholeiitic; WPB = within-plate basalt.

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scenario magmatic activity would have begun in the northern part of the CED due to lithospheric collapse at a major transform/fracture zone that results in asthenospheric upwelling. Raising the hot asthenosphere leads to extensive melting and extension, i.e. formation of a protoarc–forearc region. In the early stages of melting and extension the magmas produced are MORB-like which develop rapidly to IAT-boninitic magmas due to progressive depletion of the asthenospheric source. This refractory residue is increasingly enriched by SSZ-derived fluids, which stimulates further melting. Mantle upwelling and extension in the newly formed protoarc–forearc basin above the sinking lithosphere ends by the beginning of downdip motion of the latter (i.e. true subduction begins). This results in chilling the forearc mantle, turning it into lithosphere and migration of the magmatic activity away from the trench to form a magmatic arc. In the CED, continuous subduction probably gave birth to a back-arc spreading southward as represented by the many ophiolites in the southern part of the CED (e.g. Wadi Beririq, Gabal Ghadir and Wadi Ghadir; Fig. 1) including the investigated AMO, which display a depleted to enriched mantle source (Farahat et al., 2004). The source heterogeneity beneath back-arc basins has been explained by variable mixing of enriched and depleted mantle in addition to metasomatism by slab-derived fluids. The source of the fertile mantle is related to upwelling of asthenospheric mantle due to rifting and extension in the back-arc region (e.g. Gribble et al., 1998; Kuzmichev et al., 2005). However, Ewart and Hawkesworth (1987) attributed the supply of fertile mantle to subduction-induced mantle corner flow. Many studies (e.g. Dixon and Golombek, 1988) argued for the high crustal growth rate of the ANS using a simple arc–arc accretion model. Some authors related such high crustal growth to involvement of a pre-Pan African continent (e.g. Farahat et al., 2004; Hargrove et al., 2006), while others favoured asthenospheric uprise due to plume interaction or oblique convergence (e.g. Stein and Hofmann, 1994; Farahat, 2006; Farahat et al., 2007). According to the inferred CED geotectonic model for the CED, the raised hot asthenosphere can be accounted for by counter flow mantle convection during subduction initiation and back-arc rifting, likely induced by oblique convergence. Acknowledgements The author is greatly indebted to Profs. G. Hoinkes and C. Hauzenberger for putting some of the analytical facilities of the Institute of Earth Sciences, Graz University, Austria at his disposal. I am very grateful to Prof. R.J. Stern, an anonymous referee and the editor (Prof. N. Eby) whose critical reviews greatly improved the quality of this manuscript. References Abd El-Rahman, Y., Polat, A., Dilek, Y., Fryer, B.J., El-Sharkawy, M., Sakran, S., 2009a. Geochemistry and tectonic evolution of the Neoproterozoic Wadi Ghadir ophiolite, Eastern Desert, Egypt. Lithos 113, 158–178. Abd El-Rahman, Y., Polat, A., Dilek, Y., Fryer, B.J., El-Sharkawy, M., Sakran, S., 2009b. Geochemistry and tectonic evolution of the Neoproterozoic incipient arc–forearc crust in the Fawakhir area, Central Eastern Desert of Egypt. Precambrian Research 175, 116–134. Abdel-Karim, A.M., Azzaz, S.A., Moharem, A.F., El-Alfy, H.M., 2008. Petrological and geochemical studies on the ophiolite and island arc association of Wadi Hammariya, Central Eastern Desert, Egypt. The Arabian Journal for Science and Engineering 33, 117–138. Abdelsalam, M.G., Stern, R.J., 1996. Sutures and shear zones in the Arabian–Nubian Shield. Journal of African Earth Sciences 23, 289–310. Abu El Ela, F.F., 1985. Ophiolitic mélange of the Abu Mireiwa District, Central Eastern Desert, Egypt. Unpublished PhD Thesis, Assuit University, Egypt, 247 pp. Abu El Ela, F.F., Farahat, E.S., 2010. Neoproterozoic podiform chromitites in serpentinites of the Abu Meriewa-Hagar Dungash District, Eastern Desert, Egypt: geotectonic implications and metamorphism. The Island Arc 19, 151–164. Akaad, M.K., Abu El Ela, A.M., 2002. Geology of the basement rocks in the eastern half of the belt between latitudes 25° 30′ and 26° 30′ N central Eastern Desert, Egypt. Geological Survey of Egypt Paper No. 78, 118 pp. Ali, K.A., Stern, R.J., Manton, W.I., Kimura, J.-I., Khamis, H.A., 2009. Geochemistry, Nd isotopes and U–Pb SHRIMP zircon dating of Neoproterozoic volcanic rocks from the

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