Thermodynamic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt: Implication for fluid interaction in the Arabian–Nubian Shield ophiolites

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Journal of African Earth Sciences xxx (2014) xxx–xxx

Contents lists available at ScienceDirect

Journal of African Earth Sciences

journal homepage: www.elsevier .com/locate / ja f rearsc i

Thermodynamic modelling of Sol Hamed serpentinite, South EasternDesert of Egypt: Implication for fluid interaction in the Arabian–NubianShield ophiolites

http://dx.doi.org/10.1016/j.jafrearsci.2014.06.0011464-343X/� 2014 Elsevier Ltd. All rights reserved.

⇑ Corresponding author at: Norwegian Polar Institute, Hjalmar Johansens gt. 14,NO-9296 Tromsø, Norway.

E-mail address: [email protected] (T.S. Abu-Alam).

Please cite this article in press as: Abu-Alam, T.S., Hamdy, M.M. Thermodynamic modelling of Sol Hamed serpentinite, South Eastern Desert ofImplication for fluid interaction in the Arabian–Nubian Shield ophiolites. J. Afr. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.00

Tamer S. Abu-Alam a,b,c,⇑, Mohamed M. Hamdy a

a Geology Department, Faculty of Science, Tanta University, Tanta, Egyptb Norwegian Polar Institute, Hjalmar Johansens gt. 14, NO-9296 Tromsø, Norwayc Egyptian Institute of Geodynamic, Cairo, Egypt

a r t i c l e i n f o

Article history:Received 20 March 2013Received in revised form 1 June 2014Accepted 3 June 2014Available online xxxx

Keywords:Arabian–Nubian ShieldForearc subducted peridotiteOphiolitesCarbonatizationThermodynamic modellingT-XCO2

a b s t r a c t

The Arabian–Nubian Shield is the largest tract of juvenile continental crust of the Neoproterozoic. Thisjuvenile crust is composed of intra-oceanic island arc/back arc basin complexes and micro-continentswelded together along sutures as the Mozambique Ocean was closed. Some of these sutures are markedby ophiolite decorated linear belts. The Sol Hamed ophiolite (808 ± 14 Ma) in southeastern Egypt at theAllaqi-Heiani-Onib-Sol Hamed-Yanbu arc–arc suture represents an uncommon example of rocks thatmight be less deformed than other ophiolites in the Arabian–Nubian Shield. In order to understandfluid–rock interactions before and during arc–arc collision, petrological, mineral chemistry, whole-rockchemistry and thermodynamic studies were applied to the Sol Hamed serpentinized ophiolitic mantlefragment. These studies reveal that the protolith had a harzburgite composition that probably originatedas forearc mantle in the subducted oceanic slab. We propose that these rocks interacted with Ti-richmelts (boninite) in suprasubduction zone, which latter formed the Sol Hamed cumulates. Spinel’s Cr#associated with the whole rock V–MgO composition suggest that the harzburgites are highly refractoryresidues after partial melting up to 29%. The melt extraction mostly occurred under reducing conditions,similar to peridotites recovered from the subducted lithosphere. Protolith alteration resulted from twostages of fluid–rock interaction. The first stage occurred as a result of infiltration of concentrated CO2-richfluid released from carbonate-bearing sediments and altered basalt at the subduction zone. The alterationoccurred during isobaric cooling at a pressure of 1 kbar. The fluid composition during the isobaric coolingwas buffered by the metamorphic reactions. The second stage of fluid–rock interactions took placethrough prograde metamorphism. The increase in pressure during this stage occurred as a result ofthrusting within the oceanic crust. In this process the forearc crust was loaded by roughly 20–30 kmof overthrust rocks.

� 2014 Elsevier Ltd. All rights reserved.

1. Introduction

Arabian–Nubian Shield (ANS) in Northeast Africa and West Ara-bia is the largest tract of juvenile continental crust of the Neopro-terozoic age on Earth (Patchett and Chase, 2002; Stern et al., 2004).This crust was generated when arc terranes were created withinand around the margins of the Mozambique Ocean, which formedin association with the breakup of Rodinia �800–900 Ma (Stern,1994; Hassan et al., 2014). These crustal fragments collided asthe Mozambique Ocean closed around 600 Ma (Meert, 2003).

Due to this collision processes a supercontinent variously referredto as Greater Gondwanaland (Stern, 1994), Pannotia (Dalziel, 1997)or just Gondwana (e.g. Abu-Alam et al., 2013) was formed. The col-lision of the island arcs during ANS evolution resulted in the forma-tion of major linear suture zones of deformed ophiolitic rocksseparating less deformed volcanic arc rocks (Fig. 1).

The major suture zones of the ANS can be classified into twotypes, arc–arc and arc-continent sutures (Abdelsalam and Stern,1996). The arc–arc sutures trend mostly NE–SW representing thezones of closure of the oceanic basins between juvenile arc terr-anes at 800–700 Ma (Pallister et al., 1988; Kröner et al., 1992;Johnson et al., 2004). The Allaqi-Heiani-Onib-Sol Hamed-Yanbuand Nakasib-Bir Umq sutures are good examples of this type. Fol-lowing arc–arc collision, the ANS collided with pre-Neoproterozoic

Egypt:1

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Saudi Arabia

Al Amar

Idsas

Hulayfah

Najd Faults

B.Umq

AlWask

Thurwah

RedOnib

Wadi Halfa

Sinai

Eastern DesertEgypt

BayudaDesert

Sudan

Nakasis

Baraka

Sea

BishahHamdah

ClosureButana

QalaEn Nahl

NubaMtns Ingessana

hills

Ethiopia

Somalia

Marda F

TuluDimitri

Didessa S

Kurmuk

Akobo S

Akabo AdolaPossible shear

Nyangea S

Aswa S KarasukMoyale Mutilo F

BaragoiSekerr

MtElgon Kenya

30 40 50

0

5

10

15

20

25

Pan-African sequences(volcano-sedimentary arcassemblage, syn-,post-tectonic intrusions,gneisses and schists)

Faults and shear zones

Ophiolite Complexes

25

20

10

15

30 40

Fig. 2Gebel Ess

Gebel Uwayjah

SolHamd

Rahib

N

Fig. 1. Distribution of the ophiolites in the Arabian–Nubian Shield (modified after Vail, 1983; Hargrove et al., 2006; Ali et al., 2010; Azer et al., 2013; Abu-Alam et al.,unpublished data).

2 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

continental blocks, East- and West-Gondwana, at 750–630 Ma(Stern, 1994; Johnson et al., 2004; Abu-Alam et al., in press).Abdelsalam and Stern (1996) referred to these boundaries as arc-continent sutures. These sutures trend north–south and their bestexamples are the Nabitah suture in the east and the Keraf suture inthe west. The final collision of East- and West-Gondwana deformedthe ANS along north trending shortening zones developed between650 and 550 Ma (Abdelsalam and Stern, 1996; Stern et al., 2004;Abu-Alam and Stüwe, 2009). The well-known example of thesezones is the Hamisana shear zone, which is characterized byeast–west crustal shortening fabrics, steep folds, and thrust faults.The sutures, typically reactivated as transpressional/transcurrentzones, are located across the shield (e.g., Johnson et al., 2004, andreferences therein). Late deformation included the occurrence oflateral escape tectonics along transtensional or transpressionalsystems during the final stages of orogeny (e.g., Stern, 1994;Johnson et al., 2004; Meyer et al., 2014).

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Ophiolitic rocks are remarkably abundant in the ANS. They arescattered across most of the shield, over a distance of �3000 kmfrom the farthest north (Gebel Ess) almost to the equator, and fromRahib in the west to Gebel Uwayjah (45�E) in the east (Fig. 1). Theabundance of the ophiolites is a further indication that the Ara-bian–Nubian Shield was produced by processes similar to thoseof modern plate tectonics (Stern et al., 2004). The ophiolitic rocksof Eastern Desert (ED) of Egypt (Fig. 1) are interpreted to be formedin a suprasubduction zone (SSZ) (e.g. Ahmed et al., 2006; Azer andStern, 2007), either in forearc (Stern et al., 2004; El-Gaby, 2005;Azer and Stern, 2007; Azer et al., 2013; Hamdy et al., 2013;Khedr and Arai, 2013) or back-arc (El-Sayed et al., 1999; ElBahariya and Arai, 2003; Farahat et al., 2004). However, some stud-ies suggested that these rocks have been generated in mid-oceanridges (Ries et al., 1983; Zimmer et al., 1995; El Bahariya andArai, 2003; Farahat, 2010). The East- and West-Gondwana collisionled to obduction of the SSZ ophiolitic rocks over a continental

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

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T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 3

margin (Akaad and Abu El Ela, 2002; El-Gaby, 2005) of the West-Gondwana (Abd El-Rahman et al., 2009). Subduction was activewhile the process of ophiolitic overthrusting was operative alongthrust planes (Kröner et al., 1987; Stern, 1994). The ophiolitesare not all the same age and there are progressive changes in ageof ophiolites across the ANS (Berhe, 1990; Shackleton, 1994; Aliet al., 2010). They have an isotopic Neoproterozoic age rangingfrom 890 to 690 Ma, documenting a 200 Ma year period of oceanicmagmatism, and are caught up in �780 Ma to �680 Ma suturezones that reflect a 100 Ma year period of terrane convergence(Johnson et al., 2004; Stern et al., 2004).

In addition to the ophiolitic-, the arc-related metasedimentary,the tholeiitic and the calc-alkaline magmatic rocks (likely define anevolving, subduction related continental arc setting; Miller andDixon, 1992), other younger rocks including dikes, molasse-typesedimentary rocks, high-K volcanic rocks and alkaline graniteswere formed during the last stages of the Neoproterozoic develop-ment of the ANS. These rocks are related to a late extensional stagedeveloped in the ANS (Blasband et al., 2000; Johnson andWoldehaimanot, 2003; Ghoneim et al., 2014).

Origin of the ultramafic protolith in SSZ or in mid-ocean ridgesand the ophiolite obduction and overthrusting might cause fluid–rocks interaction and hence alteration, serpentinization and meta-somatism (e.g. Hamdy et al., 2013). Thus, much doubt existsaround origin and composition of the fluid during the alterationprocess. Some authors suggested that the alteration of theultramafic rocks – the dominant component of the ANS’s ophiolites

Serpentinite

Metagabbro

Pillow lava

Arc metavolcanic

Mélange metasedimen

Fault

22 19

36 0

5

W. D

itt

Fig. 2. Simplified geological map of Sol Hamed

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– occurred by interaction with hot fluid during seafloor weathering(e.g. Lebda, 1995; Li and Lee, 2006). Other authors believe that thealteration took place by infiltration of metamorphic and hydro-thermal fluid along major tectonic fractures during or after rockexhumation (e.g. Hyndman and Peacock, 2003; Hamdy, 2004;Hamdy and Lebda, 2007; Abu-Alam and Stüwe, 2011) or in thesubduction zone as the ultramafic rocks were a component of theforearc (Hamdy et al., 2013).

Sol Hamed ophiolite in southeastern Egypt and northeasternSudan (Fitches et al., 1983) at the Allaqi-Heiani-Onib-Sol Hamed-Yanbu arc–arc suture (Abdelsalam and Stern, 1996; Abdelsalamet al., 2003) differs from other ophiolites further north in the EDof Egypt in being an elongated and intact belt defining a near-source tectonic facies (Abdelsalam and Stern, 1996). To the north,ophiolites occur in tectonic mélanges or as olistostromal debris,indicating a distal tectonic facies. This interpretation implies theophiolitic rocks north of the Sol Hamed represent a far-travelledophiolitic nappe, transported to the north away from its corre-sponding suture. Thus the Sol Hamed rocks represent an uncom-mon example in the Eastern Desert that might be less deformedby the movement along faults that occurred after the closure ofthe Mozambique Ocean. In this work, petrological relationships,mineral chemistry, geochemistry and thermodynamic modellingare described and applied to ultramafic rocks from the Sol Hamedophiolite (Fig. 2). The results help to more clearly define the natureof the fluid–rock interactions process occurred in the intraoceaniccollision of the ANS. This is essential for understanding the genesis

s

tsThrust

0 1 2 km

22 14

36 12

W. E

l Kw

an

G.Makarib Sample locationMagnesite siteChromite site

N

area modified after Abu El Laban (2002).

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

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4 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

of the ANS Neoproterozoic ultramafic fluid–rock interactions andits relation to mineralizations.

(a)~ 5 mm

s

s

mag

maganth

anthanth

anth

anth

opxopx

opx

liz

liz

liz

liz liz

liz

2. Geological setting

Many of the ultramafic outcrops in the Arabian–Nubian Shieldare detached, scattered and isolated (Fig. 1) due to intrusion ofsyn- and post-tectonic plutons. Gass (1977) noted that these ultra-mafic bodies have tectonic contacts with other Pan-African rocks.Some of these ultramafics are recognized as ophiolites, represent-ing obducted fragments of an oceanic lithosphere that existedbetween the Proterozoic island arcs (Garson and Shalaby, 1976;El-Ramly et al., 1993). The Neoproterozoic ultramafics are repre-sented also by rare younger intrusions. The intrusive mafic–ultra-mafic complexes form undeformed small, elliptical outcrops andare commonly concentrically zoned or layered intrusion (e.g.Farahat and Helmy, 2006). Dixon (1979) estimated that the ultra-mafic bodies account for 5.3% of all Precambrian outcrops in Egypt.The Neoproterozoic ophiolites are common in the central andsouthern sectors of the Eastern Desert of Egypt (Fig. 1), where theyoccur as tectonized bodies and mélanges of pillowed metabasalt,metagabbro, and variably altered ultramafic rocks (El Sharkawyand El Bayoumi, 1979). The ultramafic rocks are mostly serpenti-nized with relicts of fresh ultramafic protolith, but sometimesinclude quartz carbonates (i.e. listwaenite), talc-carbonates, mag-nesite veins and chromite pods.

The Sol Hamed ophiolite is a part of Allaqi-Heiani-Onib-SolHamed-Yanbu arc–arc suture (Abdelsalam and Stern, 1996;Abdelsalam et al., 2003). This arc–arc suture is considered – alongwith the Ariab-Nakasib-Thurwah-Bir Umq suture farther south inArabia and Sudan (Johnson et al., 2004) to be one of the two longestand most complete Neoproterozoic ophiolite-decorated sutures inthe ANS (Azer et al., 2013). Stern et al. (1990) proposed that the All-aqi-Heiani-Onib-Sol Hamed-Yanbu suture represents a south verg-ing nappe which was refolded around a subhorizontal east–westtrending axes to produce upright antiforms and late-stage south-east verging thrusts. Vergence of the ophiolite nappe was used toinfer a north dipping subduction zone along the line of a suturewhich lies north of the Allaqi-Heiani-Onib-Sol Hamed-Yanbu ophi-olite. Ali et al. (2010) suggested two stages for the evolution of All-aqi-Heiani-Onib-Sol Hamed-Yanbu suture (�810–780 Ma and�750–730 Ma).

The ultramafic rocks of the Sol Hamed (Fig. 2) are composed ofserpentinites, chromite-bearing serpentinites and magnesite-bear-ing serpentinites, forming the base of a dismembered ophioliticsequence. The Sol Hamed ophiolite sequence comprises also

Table 1Summary of mineral assemblages of the studied ultramafic rocks.

Primary minerals OlivineOpxSpinelChromite (the inner core)

First stage of alteration and metamorphism LizarditeAntigoriteAnthophylliteMagnesiteMagnetiteChromite (inter. zone)Sulfides

Second stage of alteration and metamorphism LizarditeChrysotileTalcMagnesiteMagnetiteChromite (outer zone)

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metagabbros, pillow lavas and pelagic sediments (Abu El-Laban,2002). Serpentinites occur as both tectonized and mélange bodies.Tectonized serpentinites are massive and form ridges about 20 kmlong and about 0.4–1.8 km wide, elongated in NE–SW direction.They are bounded and thrusted over arc metavolcanics from thenorthwest. Some rock portions are extremely altered, especiallyalong thrusts and shear zones, with the development of talc, talc-carbonate and reddish brown quartz-carbonate rock (listwaenite).On the other hand at their southeastern side, Abu El Laban (2002)recorded that they are brecciated and fragmented and borderedwith a significant mélange metasediments belt (up to 1 km wide).The mélange metasediments include in addition to the serpentinitebodies metavolcanics, metagabbros, schists and rare marble. Thismélange belt separates the tectonized serpentinite bodies fromthe ophiolitic metagabbro masses. Also, the mélange metasedi-ments form low to moderate topography mainly of serpentinites,schists and marble to the east of the tectonized serpentinites and

Chlorite

Talc

Magnetite and spinel

(b)

Zoned Chromite

Chrysotile

Olivine

20 µm

Fig. 3. (a) Semi-schematic drawing showing the ophiolitic ultramafic of Sol Hamedarea, lizardite and anthophyllite are metamorphosed after orthopyroxene. Chrys-otile is metamorphosed after lizardite. Talc is after anthophyllite. Liz, opx, anth, magand s are lizardite, orthopyroxene, anthophyllite, magnesite and sulfides, respec-tively. (b) Zoned spinel in Sol Hamed serpentinite. Darker zones are richer in Cr.

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T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 5

the metagabbro. It is noteworthy that mapping of the mélange beltby Abu El Laban (2002) differs from the earlier maps of Fitcheset al. (1983) and Hussein (1981) as the contact between ultramaficand mafic rocks being sharp and marked by the presence of leuco-cratic gabbro.

Although the Sol Hamed ultramafics are completely serpenti-nized, partly serpentinized dunites in the NE part of the ultramaficbelt are found. These dunite bodies mainly contain chromititepods. In addition, the Sol Hamed serpentinites contain rare pyrox-enite bodies that can be easily distinguished from its high contentof the pyroxene relics. Due to pervasive serpentinization, fieldidentification of the Sol Hamed ultramafics as mantle or cumulatematerial was not possible – although the presence of chromitemasses in some serpentinite masses might suggest that at leastpart of the ultramafic sequence is of cumulate origin (Church,1981). Fitches et al. (1983) recorded that the Sol Hamed ophioliteincludes wehrlite and Iherzolite cumulates.

Chromitite deposits occur mainly as lenticular bodies of vari-able dimensions up to 25 m length � 6 m width, trending ENE–WSW. Thick pods are abundant in serpentinites that are mostlyderived from dunite. Micro-lenses and thin planar segregationsoccur in the serpentinized peridotite. Gradual contacts betweenmassive ore and serpentinized dunite over a meter-range are fre-quently observed. A typical contact shows gradation from fine-grained disseminated chromite in the dunite through nodular, tomassive coarse-grained chromite ore. The highly deformed chro-mite bodies are the most abundant. Magnesite veins cross-cutthe eastern periphery of the serpentinite rocks. Hamdy (2007),based on the C–O isotopes of these veins, estimated that carbonwas supplied from both geothermal fluids (giving magnesite withd13C values from –2.06‰ to –4.34‰ VPDB) and metamorphic car-bonaceous sediments (giving magnesite with d13C values from –9.44‰ to –10‰ VPDB).

Table 2Representative analyses of pyroxene and serpentine. The chemical formula of the serpentinformula was calculated based on 6 oxygen atoms.

Sample Opx Serpentine

Liz

325/1 325/2 118/2 118/3 118/4 369 333/1 333

SiO2 58.21 58.37 53.22 54.27 52.38 41.94 40.62 43.7TiO2 0.12 0.09 b.d.l b.d.l 0.06 0.03 0.12 0.08Al2O3 0.74 0.67 0.04 0.04 1.22 0.28 0.3 0.06Cr2O3 0.42 0.36 b.d.l b.d.l b.d.l 0.36 0.22 b.d.lFeO 4.85 5.09 1.10 1.12 1.10 6.21 4.2 2.66MnO 0.03 0.07 0.01 0.01 b.d.l b.d.l 0.38 b.d.lMgO 34.99 35.1 18.92 19.29 20.22 35.98 34.37 37.7CaO b.d.l 0.03 24.08 24.56 23.02 0.11 0.03 0.07Na2O b.d.l b.d.l 0.03 0.03 0.03 0.08 b.d.l 0.12K2O b.d.l b.d.l b.d.l b.d.l b.d.l 0.01 b.d.l 0.01Total 99.36 99.78 97.39 99.32 98.02 85 80.24 84.5

Si 2 2 1.974 1.974 1.927 8.162 8.283 8.36Ti 0.003 0.002 – – 0.002 0.004 0.018 0.01Al 0.03 0.027 0.002 0.002 0.053 0.064 0.072 0.01Cr 0.011 0.01 – – – 0.055 0.035 –Fe 0.139 0.146 0.029 0.031 0.028 1.011 0.716 0.42Mn 0.001 0.002 – – – – 0.066 –Mg 1.792 1.792 1.046 1.046 1.109 10.43 10.44 10.7Ca – 0.001 0.957 0.957 0.908 0.023 0.007 0.01Na – – 0.002 0.002 0.002 0.03 – 0.04K – – – – – 0.002 – 0.00Cations 3.977 3.98 4.01 4.01 4.03 19.79 19.65 19.6

di – – 0.94 0.94 0.92hed – – 0.002 0.002 0.001en 0.93 0.92 – – –fs 0.0055 0.006 – – –mgts 0.027 0.01 – – –

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3. Petrography

Variable degrees of alterations are observed in the studiedultramafic rocks. Original minerals have been preserved (Table 1)in partly altered rocks. The dominant serpentine mineral is lizar-dite, whereas chrysotile is subordinate. The lizardite forms mainlypseudomorphic mesh and bastite textures after olivine and ortho-pyroxene and sometimes occurs as interlocking and penetratinggrains (non-pseudomorphic). The abundance of textures after oliv-ine and orthopyroxene suggests harzburgite protolith. The chryso-tile occurs as cross fiber veins traversing the lizardite matrix.Serpentine minerals appear to be accompanied by shedding offine-grained magnetite, which concentrates in veins cutting zonedchromite (Fig. 3a) or along relict pyroxene cleavages. Pyroxenerelicts occur as inclusions in anthophyllite (Fig. 3a). The anthophyl-lite is a common replacement mineral of orthopyroxene, where itinitially grows along cleavage planes and eventually replaces thewhole grain. Talc is not abundant in the studied serpentinites. Itforms fine shreds, dense fibers and medium grained flaky crystals(0.01–0.04 mm). Perfect cleavage, straight extinction and highinterference colors are characteristic features of the talc. The talcis pseudomorphic after anthophyllite. It is homogenous and com-monly associated with the alteration of orthopyroxene. All ser-pentinite samples contain zoned-chromite (Fig. 3b) and sulfidegrains. Chromite occurs as disseminated subhedral and anhedralcrystals of reddish brown color. Some chromite grains look homo-geneous in reflected light. The carbonate minerals are opticallynegative and show perfect cleavage parallel to the crystal face(1011) with colorless to yellowish brown color. The colored crys-tals may indicate Mg-rich and Fe-rich components (e.g. magnesiteand siderite). The spinel minerals are opaque crystals with nocleavage. The spinel crystals show zonation where the cores aredarker than the rims (Fig. 3b).

e minerals was calculated based on 28 oxygen atoms and ignoring the H2O, pyroxene

Chr

/2 347/1 347/2 278/1 310 278/2 369 325/1 325/2

8 41.3 41.56 43.04 43.96 44.54 43.23 42.6 42.47b.d.l 0.01 0.15 0.08 0.07 b.d.l 0.01 b.d.l0.26 0.14 0.14 0.14 0.28 0.21 b.d.l 0.130.16 0.09 0.34 0.12 b.d.l 0.15 b.d.l 0.023.36 3.99 1.85 0.94 1.37 2.11 2.05 1.810.22 0.17 0.09 b.d.l 0.01 0.01 b.d.l b.d.l

4 38.68 38.06 36.88 38.14 37.78 37.17 36.68 38.42b.d.l b.d.l 0.13 b.d.l 0.06 0.18 0.19 b.d.lb.d.l b.d.l 0.01 0.1 0.01 0.09 b.d.l 0.02b.d.l 0.01 0.03 0.08 0.02 0.08 0.03 0.01

2 83.98 84.03 82.66 83.56 84.14 83.23 81.56 82.88

9 8.036 8.097 8.385 8.418 8.47 8.375 8.41 8.2582 – 0.001 0.022 0.012 0.01 – 0.001 –4 0.06 0.032 0.032 0.032 0.063 0.048 – 0.03

0.025 0.014 0.052 0.018 – 0.023 – 0.0035 0.547 0.65 0.301 0.151 0.218 0.342 0.338 0.294

0.036 0.028 0.015 – 0.002 0.002 – –5 11.21 11.05 10.71 10.88 10.71 10.73 10.79 11.134 – – 0.027 – 0.012 0.037 0.04 –4 – – 0.004 0.037 0.004 0.034 – 0.0082 – 0.002 0.007 0.02 0.005 0.02 0.008 0.0024 19.92 19.88 19.56 19.58 19.49 19.62 19.59 19.73

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

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90

9090

90

Mg

Fe+Mn Al+Cr

Mg

(a)

50 0 50

Fe

Cr

Granulite-faciesSpinels

Upper-amphibolite-facies spinels

Lower-amphibolite-facies spinels

Fe-spinels

facies Cr-spinels

Al

3+

Greenschist-

3+3+(d)Core intermediate zone rim

a bc

de

f

I

0.0

0.2

0.4

0.6

0.8

1.0

1.0 0.8 0.6 0.4 0.2 0.0

Mg# (Spl)

Cr#

(Spl

)

III

II

(b)

50 0 50

FeO

CaO MgO

(c)

Fig. 4. Mineral chemistry. (a) Substitution in serpentine. Aluminum and chromium are grouped together, as they tend to vary sympathetically. (b) Mg# vs. Cr# variationdiagram for the investigated spinel. field I represents Cr-spinels of mantle peridotite, field II represents magnetite from metamorphic rocks and field III is magnetite fromunmetamorphosed igneous rocks (fields after Roeder, 1994; Mondal et al., 2001), (c) MgO–CaO–FeO ternary diagram showing the chemical composition of the carbonateminerals, (d) compositional changes in spinels expressed in a triangular Cr–Fe3+–Al3+ plot with reference to the fields of spinel types: a – chromian magnetite, b – aluminianmagnetite, c – ferrian spinel, d – chromian spinel, e – ferrian chromite and f – aluminian chromite (Stevens, 1944) and the different metamorphic facies defined by Purvis et al.(1972), Evans and Frost (1975) and Suita and Strieder (1996).

6 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

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T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 7

4. Mineral chemistry

Different mineral phases were examined in the Institute of Geo-logical Sciences of Polish Academy of Sciences (IGS-PAS). The min-eral analyses were carried out by JEOL-JXA-840A scanning electronmicroscope equipped with Link Analytical AN-1000/855 energydispersive X-ray spectrometer. The analytical conditions were15 kV accelerating voltage and 35 nA beam current. The followingstandards were used during the measurements: enstatite for MgOand SiO2; adularia for K2O and Al2O3; ferrosilite for FeO; titanite forTiO2 and CaO; jadeite for Na2O; spessartine for MnO; chromite forthe Cr2O3. Mineral formula and activity of the end-members werecalculated by AX program (http://www.esc.cam.ac.uk/research/research-groups/holland/ax). The mineral abbreviations which willbe used in the following sections are from Holland and Powell(2011).

Table 3Representative analyses of talc. The chemical formula and the end-members activitieswere calculated by the AX program and based on 11 oxygen atoms and ignoring theH2O.

Sample Talc

306/1 306/2 306/3 306/4 306/5

SiO2 63.30 62.11 61.65 63.08 61.09TiO2 0.05 b.d.l 0.04 0.07 0.22Al2O3 0.03 0.14 0.28 0.05 b.d.lCr2O3 0.02 0.14 0.26 0.03 b.d.lFeO 3.43 3.73 3.90 2.85 3.11MnO b.d.l b.d.l 0.04 b.d.l b.d.lMgO 28.13 27.84 27.17 28.57 27.60CaO 0.09 0.02 b.d.l 0.16 0.16Na2O 0.13 b.d.l 0.16 0.18 0.13Totals 95.20 94.00 93.50 95.00 92.31

Si 4.053 4.036 4.034 4.041 4.035Ti 0.002 – 0.002 0.004 0.011Al 0.002 0.011 0.022 0.004 –Cr 0.001 0.007 0.014 0.002 –Fe3+ – – – – –Fe2+ 0.184 0.203 0.213 0.153 0.172Mn – – 0.002 – –Mg 2.685 2.696 2.65 2.728 2.717Ca 0.006 0.001 – 0.011 0.011Na 0.017 – 0.021 0.022 0.017Sum 6.951 6.955 6.957 6.964 6.962

ta 0.71 0.72 0.68 0.75 0.74fta 0.00023 0.00031 0.00036 0.00013 0.00019

Table 4Carbonate analyses. The chemical formula was calculated based on 2 cations.

Sample Carbonate minerals

300/1 300/2 300/3 306/1 306/2 306/3

SiO2 0.19 0.15 0.04 0.11 0.23 0.18Cr2O3 0.11 0.04 0.09 b.d.l 0.07 b.d.lFeO 13.47 11.17 13.42 14.1 12.33 8.46MnO 0.09 0.25 0.13 0.33 0.33 0.22MgO 36.19 38.06 36.82 35.53 36.53 40.14CaO 0.11 0.23 0.17 0.15 0.15 0.24Totals 50.16 49.9 50.67 50.22 49.64 49.24

Si 0.006 0.004 0.001 0.003 0.007 0.005Cr 0.003 0.001 0.002 – 0.002 –Fe2+ 0.343 0.28 0.338 0.361 0.315 0.21Mn 0.002 0.006 0.003 0.009 0.009 0.006Mg 1.643 1.701 1.65 1.622 1.663 1.772Ca 0.004 0.007 0.005 0.005 0.005 0.008Sum 2 2 2 2 2 2

mag 0.84 0.86 0.84 0.83 0.84 0.89sid 0.26 0.21 0.25 0.27 0.24 0.17

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CaO content is below 0.03 wt% in the orthopyroxene and FeOcontent is in the range of 4.85–5.09 wt% while MgO content isaround 35 wt% (Table 2). This reveals that the main orthopyroxeneend-member is enstatite. While diopside is the main end-memberfor the clinopyrocene. SiO2 content of the serpentine rangesbetween 40.62 and 44.54 (Table 2). Al2O3 is in the range of belowthe detection limit up to 1.79 wt%. The serpentines are classified tolizardite and chrysotile (Table 2) based on the FeO content(Norman, 1968), where the chrysotile contains lower FeO (0.94–2.11 wt%) than the lizardite (2.66–6.21 wt%). MgO ranges between34.37 to 38.68 wt%. The MgO and the FeO ranges indicate ionicsubstitution between Fe2+ and Mg2+ (Fig. 4a). FeO and Cr2O3 con-tents in lizardite increase (0.94–6.21 wt% and from below detec-tion limit to 0.36 wt% for FeO and Cr2O3, respectively) distinctlywith increasing degree of alteration from partly to completely ser-pentinized rocks (Table 2). Chrysotile shows that Al and Cr are rel-atively immobile during recrystallization of lizardite and thereforeremain in their original crystal lattice.

Low Al2O3 and TiO2 contents in talc chemistry reveal limitationin substitution between Si, Ti and Al (Table 3). The activities of talcand Fe-talc end-members are in the range of 0.68–0.85 and0.00013–0.00061, respectively. Table 4 shows chemical analysesof the carbonate minerals. The high concentrations of MgO andFeO (35.53–40.14 and 8.46–14.1 wt%, respectively) indicate highactivity of the magnesite and the siderite end-members (Fig. 4c).The CaO content is in the range of 0.04–0.27 wt% revealing lowactivity of the calcite.

Three compositional zones are distinguished for the spinel min-erals. Compositions of core, intermediate and rim zones are givenin Table 5 and plotted in Al–Cr–Fe3+ triangle of Stevens (1944)(Fig. 4d). Cores and intermediate zones have aluminian chromiteto ferritchromite composition. Composition of the outer rim isCr-magnetite which is nearly devoid of Al and lie along the Cr–Fe3+ sideline (Fig. 4d). All the core compositions have Cr# andMg# displayed in the mantle chromite field of Roeder (1994)(Fig. 4b).

The variation in the spinel composition can be interpreted as aresult of chemical alteration under hydrothermal conditions(Abzalov, 1998; Barnes, 2000; Proenza et al., 2004). The alterationis accompanied by decrease in Al, Mg and Cr contents and conse-quently increase in Fe3+ and Fe2+. Apparently with the increasingof the alteration, Fe releases from olivine and orthopyroxene andCr releases from chromite and are accommodated in the serpen-tines. According to Barnes (2000), the Fe3+-rich aluminian chro-mite–ferritchromite zone were formed by reactions between the

306/4 306/5 306/6 368/1 368/2 368/3 368/4

0.05 0.04 0.09 0.13 0.16 0.1 0.09b.d.l b.d.l b.d.l 0.05 0.04 0.02 0.0713.03 13.08 13.79 8.97 13.06 13.95 10.630.03 0.08 0.17 0.24 0.25 0.25 0.2736.57 36.09 36.63 39.53 36.58 36.08 38.070.04 0.21 0.27 0.17 0.21 0.21 0.1549.72 49.5 50.95 49.09 50.3 50.61 49.28

0.002 0.001 0.003 0.004 0.005 0.003 0.003– – – 0.001 0.001 0 0.0020.333 0.336 0.346 0.224 0.331 0.354 0.2690.001 0.002 0.004 0.006 0.006 0.006 0.0071.664 1.653 1.638 1.759 1.65 1.63 1.7150.001 0.007 0.009 0.005 0.007 0.007 0.0052 2 2 2 2 2 2

0.85 0.84 0.83 0.89 0.84 0.83 0.870.25 0.25 0.26 0.18 0.25 0.26 0.21

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

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Table 5Spinel group analyses. The chemical formula was calculated based on 24 oxygen atoms. b.d.l is below detection limit.

Spinel

325/core2

325/rim2

325/core5

325/rim5

118/rim1

118/core1

118/co-ri1

333/rim1

333/core1

333/co-ri1

333/rim2

333/core2

333/co-ri2

310/rim2

310/core2

310/co-ri2

SiO2 0.07 0.05 0.07 0.23 0.13 0.24 b.d.l 0.21 0.05 0.02 0.22 0.19 0.12 0.21 0.22 0.04TiO2 0.36 0.23 0.21 0.18 b.d.l 0.22 0.27 b.d.l 0.15 b.d.l 0.02 0.21 0.13 b.d.l 0.08 0.11Al2O3 8.81 9.23 8.28 8.4 b.d.l 6.27 6.32 0.12 5.4 6.53 0.1 6.67 6.5 b.d.l 5.57 6.49FeO 34.12 34.47 31.23 32.51 89.78 32.88 31.47 90.89 34.12 32.71 91.34 31.6 32.14 90.03 30.39 30.7Cr2O3 51.38 51.1 53.76 52.76 1.64 56.51 56.62 1.56 56.25 55.26 1.34 55.82 55 2.1 56.98 56.84MnO 0.39 b.d.l 0.36 0.51 0.46 0.18 0.92 0.5 0.02 0.54 0.41 0.66 1.09 0.45 0.85 0.25MgO 4.23 4.03 4.96 4.56 1 4.03 4.52 0.88 3.67 4.4 1.05 4.07 4.06 0.75 4.53 5.17NiO b.d.l 0.72 0.48 0.17 0.35 b.d.l 0.22 0.72 0.01 b.d.l 0.23 0.02 0.25 0.38 0.34 b.d.lTotal 99.36 99.83 99.35 99.32 93.36 100.33 100.34 94.88 99.67 99.46 94.71 99.24 99.29 93.92 98.96 99.6

Si 0.02 0.014 0.02 0.065 0.052 0.068 – 0.083 0.014 0.006 0.087 0.054 0.035 0.084 0.063 0.011Al 2.96 3.084 2.76 2.809 – 2.099 2.113 0.056 1.84 2.207 0.047 2.249 2.203 – 1.889 2.17Ti 0.08 0.049 0.045 0.038 – 0.047 0.058 – 0.033 – 0.006 0.045 0.028 – 0.017 0.023Cr 11.56 11.451 12.018 11.83 0.522 12.687 12.692 0.488 12.852 12.526 0.42 12.623 12.499 0.663 12.956 12.744Mn 0.094 – 0.086 0.123 0.157 0.043 0.221 0.168 0.005 0.131 0.138 0.16 0.266 0.152 0.207 0.06Mg 1.796 1.705 2.093 1.93 0.6 1.708 1.913 0.52 1.583 1.883 0.621 1.737 1.742 0.447 1.944 2.188Ni – 0.16 0.11 0.04 0.11 – 0.05 0.23 – – 0.07 – 0.06 0.12 0.08 –Fe2+ 6.11 6.135 5.711 5.907 7.133 6.249 5.816 7.082 6.412 5.986 7.171 6.103 5.932 7.281 5.769 5.752Fe3+ 1.39 1.402 1.157 1.258 15.426 1.099 1.137 15.373 1.261 1.261 15.44 1.029 1.235 15.253 1.075 1.052Mg# 0.227 0.217 0.268 0.246 0.079 0.214 0.247 0.068 0.197 0.239 0.079 0.221 0.227 0.057 0.252 0.275Cr# 0.726 0.718 0.754 0.744 0.0321 0.798 0.796 0.030 0.805 0.783 0.026 0.793 0.784 0.041 0.813 0.798Fe3+# 0.087 0.087 0.072 0.079 0.967 0.069 0.071 0.965 0.079 0.078 0.970 0.064 0.077 0.958 0.067 0.065

mt 0.004 0.006 0.005 0.004 0.90 0.0013 0.0015 0.90 – 0.0009 0.89 0.0026 0.0012 0.90 0.0008 0.0018cmt 0.39 0.39 0.41 0.40 0.00058 0.48 0.46 0.00052 0.50 0.45 0.0004 0.47 0.46 0.001 0.48 0.45

8 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

aluminian chromite grains and surrounding magnetite rims, whereits size increases by increasing grade of metamorphism.

The studied spinels show metamorphic conditions correspond-ing to that of the upper greenshist to the transitionalngreenschist–amphibolite facies (Fig. 4d). The metamorphism of the Sol Hamedophiolitic peridotites is similar to that determined for many meta-morphosed ultramafic rocks in the Eastern Desert (e.g. Ghoneimand Aly, 1986; Azer and Khalil, 2005; Hamdy and Lebda, 2007).Ultramafic bodies in orogenic belts are commonly metamorphosedto the same grade as the surrounding rocks (Winter, 2001) and aredescribed by Evans (2004) as ‘‘isofacial’’. The metamorphic faciesthat prevailed during the evolution of the Egyptian Neoproterozoic

Table 6Representative whole-rock chemistry of Sol Hamed serpentinites. Major oxides are in wt%

Sample 306 310 323 324 325

Major oxides and LOI (wt%)TiO2 0.06 0.06 0.01 0.04 0.02SiO2 41.49 41.49 39.14 38.60 39.76Al2O3 0.48 0.88 0.42 0.36 0.39Fe2O3 6.79 7.49 6.58 7.99 7.84MnO 0.09 0.07 0.09 0.07 0.08MgO 38.29 39.11 39.22 38.88 38.37CaO 0.18 0.12 0.13 0.24 0.17Na2O 0.09 0.17 0.03 0.01 0.01K2O 0.01 0.05 0.01 0.00 0.01Sum 87.49 85.94 85.61 86.15 86.63LOI 10.60 13.14 13.49 13.20 13.07

Trace elements (ppm)Cr 2742.20 2696.98 2706.00 2652.89 2677.8Co 166.79 120.41 119.56 162.00 164.55Ni 2381.21 1799.16 2055.34 2055.90 2072.5Cu 63.83 14.97 23.89 51.52 42.15Zn 57.11 13.32 25.68 23.09 33.25Sr 55.04 46.61 48.12 61.96 59.05V 40.31 29.83 27.41 37.11 25.98Ba 35.24 15.00 24.98 15.14 19.88Pb 4.70 13.15 15.75 14.96 17.32Cd 2.00 2.00 3.31 3.30 2.13Li 10.08 4.99 6.99 6.15 5.11Rb 1.80 0.97 1.34 1.51 1.44

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basement complex was up to the transitional greenschist–amphib-olite facies (Azer and Khalil, 2005).

5. Whole-rock chemistry

Representative bulk rock chemistry of the Sol Hamed serpenti-nite is given in Table 6. Chemical analyses of major and some traceelements were carried out at the geochemistry laboratory of theIGS-PAS. Concentrations of major and trace elements were deter-mined after microwave-assisted acid digestion with atomicabsorption spectrophotometer (AAS-PU 9100xUNICAM). Before

, trace elements are in ppm. b.d.l is below detection limit.

333 340 347 369 378

0.02 0.02 0.02 0.01 0.0339.54 39.30 38.41 38.55 41.260.20 0.91 0.31 0.80 0.148.51 8.24 7.54 8.70 5.990.12 0.08 0.09 0.09 0.0338.73 39.39 38.43 39.83 39.350.05 0.18 0.14 0.65 0.040.00 0.22 b.d.l 0.01 0.010.01 0.03 b.d.l 0.01 b.d.l87.17 88.37 86.00 88.65 86.8211.27 11.46 12.99 10.27 12.81

0 2582.89 2699.83 2717.47 2703.71 2420.37150.42 156.28 154.23 116.20 135.18

2 1836.13 1815.06 2060.51 1650.08 1996.1353.09 54.68 8.27 46.32 32.9718.15 37.80 16.55 11.98 23.0672.02 49.89 89.11 48.08 47.9322.15 39.77 14.00 33.93 19.6617.06 16.34 45.00 20.00 29.8819.88 17.00 4.63 24.04 14.402.11 3.31 3.31 3.43 2.887.96 9.00 9.93 1.68 7.141.65 0.32 0.86 0.17 1.27

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T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 9

digestion samples were heated to 1100 �C to determine loss onignition (LOI). Analytical precision was better than 0.5% for majorelements and 4 ppm for trace elements.

Due to the almost complete serpentinization of some of the SolHamed peridotites, modal compositions could not be determined.Therefore, normative compositions were calculated from anhy-drous analyses using the CIPW norm, assuming a Fe2O3/FeO ratioof 0.2 (Melcher et al., 2002). The normative contents of olivine,orthopyroxene, and clinopyroxene of the studied Sol Hamed serp-entinites classify them as harzburgites. The authors are well awareof the fact that serpentinization may severally change the originalsilica, magnesium and calcium values. In addition, Al, Na and Ksubstitute Si, Fe, Mg and Ca in primary mantle silicates (especiallyin clinopyroxene) in variable amounts, so that the calculated nor-mative diopside values are only a minimum estimate of the pri-mary clinopyroxene. As a result, rock composition could shiftfrom the lherzolite into the harzburgite. However, trace elementvalues in the Sol Hamed peridotites are typical of residual mantle(e.g. high Cr (2.696–2.742 ppm), Ni (1.650–2.381 ppm) and Co(116.20–166.79 ppm)), consistent with the harzburgite protolith.In contrast, the contents of Ba, Pb, Sr and, Li are highly concen-trated relative to depleted and pristine mantle peridotites(McDonough and Sun, 1995). This enrichment in the fluid-mobileelements may be directly related to the serpentinization processor due to metasomatism by subduction-related fluids (Hamdyet al., 2013).

6. Discussion

6.1. Origin and tectonic setting of the serpentinite protolith

Earth contains two main shallow mantle domains: sub-oceaniclithosphere and sub-continental lithosphere. The Sol Hamedharzburgite falls within the oceanic array (Niu, 2004) in MgO/SiO2–Al2O3/SiO2 space (Fig. 5). The oceanic array is parallel to the

DMM

PM

0.5

0.6

0.7

0.8

0.9

1.0

0.0 0.05 0.10 0.15

1.1

Melt infiltrationSea-floor

weathering

Terrestrial melting array

Al2O3/SiO2

MgO

/SiO

2

Fig. 5. Whole rock MgO/SiO2–Al2O3/SiO2 plot. The terrestrial array is a compilationof subcontinental peridotites (Hart and Zindler, 1986) and represents a meltdepletion trend. The Sol Hamed serpentinites plot offset to lower MgO/SiO2 valuesbecause of alteration. Compositions of depleted MORB mantle (DMM; Workmanand Hart, 2005), primitive mantle (PM; McDonough and Sun, 1995) and seafloorweathering trend (Snow and Dick, 1995) are plotted for comparison.

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terrestrial array but offset to lower MgO/SiO2 values, presumablydue to loss of MgO during low-temperature seafloor weatheringand not due to the serpentinization process itself (Snow andDick, 1995; Niu, 2004). Oceanic peridotites may originate in a vari-ety of tectonic environments including mid-ocean ridge (MOR),suprasubduction zone (SSZ) and rifted margins settings. We termthese suprasubduction zone (SSZ) peridotites (Pearce et al.,1984); a group that incorporates peridotites from both island arcsand spreading centers above subduction zones. These discretegenetic types are distinct in mineralogical and geochemical charac-teristics of mantle residues. Composition of the unaltered acces-sory spinel is extensively used as a petrogenetic and geotectonicindicator (e.g. Barnes and Roeder, 2001). Chromium numbers [Cr/(Cr + Fe3+ + Al)] higher than 0.6 are usually restricted to subduc-tion-related rocks (Dick and Bullen, 1984). Ishii et al. (1992) usedthe Mg# [Mg/(Mg + Fe2+)] and Cr# of the spinel to discriminatebetween peridotites from MOR, forearc and back-arc settings. Spi-nels from the Sol Hamed serpentinites lie in the chemical space ofthe forearc peridotite (Fig. 6a) and distinctly higher than spinelsfrom MOR and back-arc basin in the Cr#. This indicates that theSol Hamed serpentinites represent a fragment of oceanic litho-sphere that has been incorporated above subduction zone in a fore-arc. This interpretation is consistent with the tectonic settingproposed by Church (1988) for the Sol Hamed ophiolite.

Despite this forearc nature of the Sol Hamed peridotites, TiO2

contents of the studied spinel are plotted in the fields of the com-mon area between the SSZ peridotite and the high-Ti arc (Fig. 6b)and between the forearc peridotite and the higher Ti-boninite field(Fig. 6c). We propose that this higher TiO2 contents at a given C#and Al2O3 content in spinels may be due to interaction with Ti-richmelts (boninite) in SSZ. These melts might be those which formedcumulates as a part of the ultramafic sequence at the Sol Hamed.Melt-rock reaction produced boninitic melt and porous duniticchannels in which the mixing/mingling of melts promotes crystal-lization of monomineralic high-Cr chromian spinel (González-Jiménez et al., 2011; Hamdy and Lebda, 2011) of the Sol Hamedchromitite deposit. According to the melt–rock interaction modeland despite the controversy concerning the importance of waterin the formation of chromitites, the Cr# of the spinel is controlledby both the degree of depletion of the mantle source, due to previ-ous melting, and the degree of the second-stage melting. The latteris presumably controlled mainly by the melt/rock ratio, togetherwith temperature and compositions of the melt.

It is noteworthy that the Sol Hamed serpentinites are plottedoutside the field of spinels in central Eastern Desert serpentinites,to a lower Mg# (Fig. 6a). This implies for a possible different chro-mian spinel in peridotites that locate to the extremely southernEastern Desert of Egypt. The low Mg# at a given Cr# of spinel isdue to low equilibration temperature (Khedr and Arai, 2013). Com-positions of the chromian spinel in the Sol Hamed serpentinites arenot similar with those of the chromian spinel in Arais serpentinitesof the south Eastern Desert of Egypt (Khedr and Arai, 2013). Thelatter, in contrast, are similar to the compositions of the chromianspinel in the serpentinites of the central Eastern Desert of Egypt.

However, Church (1988) based on the presence of wehrlite andIherzolite cumulates and the absence of troctolitic cumulates sug-gested that the Sol Hamed ophiolite represents slices of primitivesuprasubduction zone that were slit by activity of strike-slip faults.The association of forearc cumulate rocks that crystallized insequence of olivine–(chromite; high Cr)–clinopyroxene–orthopy-roxene–plagioclase and boninitic volcanic rocks that crystallizedin sequence of olivine–orthopyroxene–clinopyroxene–plagioclasewith a ductile fault zone in the Troodos ophiolite (Murton, 1986)and with a spreading center in the case of the Betts Cove ophioliteof Newfoundland (Coish and Chrch, 1979; Coish et al., 1982) mightsuggest that ophiolites with these characteristics owe their

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Boninite

Back-arc

Mid-Ocean Ridge

100Mg/Mg+Fe

100C

r/Cr+

Al

20406080

20

40

60

80

Forearc

100

00100

0 10 20 30 40 500.01

0.1

1

10

Al2O3 wt%

TiO

2 w

t%

MOR-typeperidotite

archigh-Ti

Suprasubductionzone peridotite

MORB

arc

low

-Ti

Cr#

TiO

2 w

t%

0 0.2 0.4 0.6 0.8 1.00

0.1

0.2

0.3

0.4

MOR

Forearcperidotite

Boninite

(a)

(b) (c)

Fig. 6. Composition of spinels compared with those in modern peridotites. (a) Data are plotted on 100Cr/Cr + Al (Cr#) vs. 100 Mg/Mg + Fe (Mg#) diagram, modified after Dickand Bullen (1984). The low Mg# at a given Cr# of spinels is due to low equilibration T. The fields are after Bloomer et al. (1995). Data of chromian spinels in serpentinites fromthe Central Eastern Desert of Egypt (Serp in CED) are obtained from the literature (Ahmed et al., 2001; Azer and Khalil, 2005; Khalil and Azer, 2007; Farahat, 2008; Farahatet al., 2011). (b) TiO2 vs. Al2O3 (Kamenetsky et al., 2001). (c) Cr# vs. TiO2. (Barnes and Roeder, 2001). MORB, mid-ocean ridge basalt.

10 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

preservation in part to their origin as strike-slip fault sliversdetached from the frontal part of arcs as a result of obliquesubduction.

Yet, the forearc affinity of the studied serpentinized peridotitesimposes more debates on whether they were formed in mantlewedge beneath the overriding plate or in subducting slab. Recently,Deschamps et al. (2013), based on the compilation of �900 geo-chemical data of abyssal, mantle wedge and exhumed serpenti-nized peridotites after subduction, could discriminate betweenforearc serpentinites from different settings. Based on REE patterns

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and Ti content of these rocks, Deschamps et al. (2013) estimatedthe nature of the initial protolith for serpentinites, as well as thegeological settings in which they were formed. The studied SolHamed peridotites have Ti content analogous to that of the sub-ducted slab serpentinites. Hellebrand et al. (2001) tested whichtrace elements correlate with major element indicators of partialmelting in central Indian ridge peridotites. The most common ofthese is the Cr# in spinel. They found a well-defined correlationbetween moderately incompatible elements, such as heavy rareearth elements (HREEs) in clinopyroxene with spinel Cr#.

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T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 11

Hellebrand et al. (2001) developed an empirical equation (F = 10 ln(Cr#) + 24) to estimate the degree of melting F (in percent) as afunction of spinel Cr#. Using the equation of Hellebrand et al.(2001), the estimated melting in the studied peridotites rangesfrom 20% to 22%. Pearce and Parkinson (1993) have shown thatmodelling of partial melting processes is best achieved with ele-ments unaffected by metasomatism, such as Ni and Co (compati-ble), Sc, V, Ga, Al (slightly to moderately incompatible), Y, Ti, andHREE (incompatible). Using the whole-rock compositions of Vand MgO (anhydrous), the Sol Hamed peridotites are comparablewith those of Lee et al. (2003) and underwent melt extraction�26–29%. High melting degree of the Sol Hamed mantle reservoiris in accordance with its formation in the forearc. Bonatti andMichael (1989) proposed that mantle melting ranges from nearlyzero for undepleted continental peridotites to about 10–15% melt-ing for rifted margins to 10–25% melting associated with mid-ocean ridge (MOR) peridotites to 30% for peridotites recoveredfrom forearcs, which generally form during the early stages inthe evolution of the associated subduction zone (Bloomer et al.,1995). In contrast to the harzburgites that represent residual man-tle after extensive melting, the dunites and wehrlites in the SolHamed ophiolite reflect melt–wallrock interactions (Stern et al.,2004).

Serpentinization plays a role in redox conditions of the mantlewhich change the oxidation state of redox-sensitive elements. Itis recognized that, for a constant amount of Fe remaining, theratios of ((Fe2O3 � 1000)/(Fe2O3 + FeO)) in bulk rock increase withthe degree of serpentinization. To determine redox conditions for

H2O

di t

r fa

H2O

tr fa

tr fs

H2O

chr

hed

atg

ta

hed

atg

hed

fa ta

100 150 200 250 300 350 400 4200

700

1200

1700

2200

2700

3200

3700

4200

4700

5200

5700

6200

1) tr fo = di en H2O

5) clin = sp fo anth H O 2

3) clin = sp fo en H O 2

4) fo anth = en H O 2

7) clin fa = herc fo anth H O 2

6) tr fa = di en fs H2O

2) fa anth = fs en H2O

8) sp fo ta = clin anth

Reaction equations are written with the high T assemblage to the right ofthe equal sign

15) clin ta = H O sp anth 2

16) clin anth fa = herc en H O 2

14) di clin fa = herc tr fo H O 2

9) fa ta = fs anth H O 2

10) clin fa ta = anth herc H O 2

12) clin anth = sp en H O 2

13) clin fa = herc fo en H O 2

11) tr hed fa = di fs H O 2

CFMASH

17) di clin = sp tr fo H O 2

liz =

br a

tg (

Evan

, 200

4)

clin

hed

ta liz

= c

hr (C

hern

osky

, 197

5)

21) di mgts clin = H2O tr sp

26) herc ta hed = H2O mgts tr fs

29) herc fta ta = H2O mgts fs

24) mgts clin = H2O ta sp

30) ta sp = H2O mgts anth

23) fta clin = H2O herc ta fs

20) fa fta = H2O fs

18) fta atg = H2O ta fa

22) hed br clin = H2O di herc fa

28) clin br = H2O sp fo

19) hed br atg = H2O di fa

27) br fa clin = herc fo H2O

25) br atg = H2O fo

18

19

20

21

22

23

H2O

fo a

tg

clin en

sp fo ta

P(bar)

Fig. 7. P–T grid in the system CFMASH for atg, chr, en, fs, di, hed, fo, fa, anth, tr, clin, treactions liz = br + atg, liz = chr, liz = ta + fo + clin + H2O are used here after Evans (20abbreviations are after Holland and Powell (2011).

Please cite this article in press as: Abu-Alam, T.S., Hamdy, M.M. ThermodynaImplication for fluid interaction in the Arabian–Nubian Shield ophiolites. J. Afr

the protoliths of serpentinites formed after the Neoproterozoicperidotites of study, we used the V–MgO composition proposedby Lee et al. (2003), since we know that V records fO2 during man-tle melting (e.g. Lee et al., 2003; Pearce and Parkinson, 1993). TheSol Hamed peridotites have compositions between QFM and QFM-1 (QFM-1; this refers to log fO2 (QFM) = log units relative to quartz–fayalite–magnetite buffer). This estimates that melt extraction inmost peridotites was reducing conditions. This is in agreementwith the serpentinites from the subduction zone (Parkinson andArculus, 1999).

6.2. Thermodynamic modelling

All the thermodynamic calculations in the following sectionswere calculated by THERMOCALC (Powell and Holland, 1988), Per-Ple_X (Connolly, 1990) and using the internally consistent datasetof Holland and Powell (2011). Lizardite bearing reactions whichwere proved experimentally (i.e. liz = br + atg (Evans, 2004),liz = chr (Chernosky, 1975), liz = ta + fo + clin + H2O (Caruso andChernosky, 1979)) will be only used (Fig. 7).

Fig. 7 shows a P–T grid in the system CFMASH for the followingend-members: atg, chr, en, fs, di, hed, fo, fa, anth, tr, clin, ta, sp,herc, mgts, fta, br, H2O. Activity of the H2O is imposed to be theunity therefore all the CO2 bearing phases are not seen in this grid.The P–T grid shows forty-six univariant equilibria, five invariantpoints and three experimental lizardite bearing reactions. All theH2O bearing univariant reactions show steep slope in the P–Tspace. Consequently these reactions can be used as temperature

anth

6 8

sp fo anth

herc fo ta

ant

h H

2O

H2O

her

c tr

fs

tr fo

H2O

her

c tr

fa H

2O

her

c fa

ta H

2O

clin en fa

clin

hed

fs

di a

tg

en ta

fo ta

clin en

3

4 5

7

clin

hed

fo ta

H2O

at

g

H2O

her

c di

tr fa

herc fo ta

1 2

9

10

11

12

13

14

15

16

50 500 550 600 650 700 750

clin anth fa

4

17

liz =

ta fo

clin

H2O

(Car

uso

& C

hern

osky

, 197

9)

herc fo anth clin en fa

fs c

lin

25

24

27

28

29

26

30

[atg chr fs di hedfa tr clin sp herc

fta br mgts]

[atg chr fs di hed trta sp fta br mgts]

[atg chr fs di hed fatr ta herc fta br mgts]

[atg chr fs di hed trsp H2O fta br mgts]

[atg chr fs di herc hed trfa H2O fta br mgts]

T CO

a, sp, herc, mgts, fta, br, H2O. Activity of the H2O is imposed to be the unity. Note:04), Chernosky (1975) and Caruso and Chernosky (1979), respectively. Mineral

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

Page 12

anth

8

sp fo anth

herc fo ta

ant

h H

2O

clin en fa

en ta

fo ta

clin en

herc fo ta

2

9

10

12

15

16

clin anth fa

4

liz =

br a

tg (

Evan

, 200

4)

liz =

ta fo

clin

H2O

(Car

uso

& C

hern

osky

, 197

9)herc fo anth clin en fa

30

[atg chr fs di hedfa tr clin sp herc

fta br mgts]

[atg chr fs di hed trta sp fta br mgts]

[atg chr fs di hed fatr ta herc fta br mgts]

[atg chr fs di hed trsp H2O fta br mgts]

clin en

sp fo ta 100 200 300 400 500 600 650 700 750

200

700

1200

1700

2200

2700

3200

3700

4200

4700

5200

5700

6200

4) fo anth = en H O 2

2) fa anth = fs en H2O

8) sp fo ta = clin anth

Reaction equations are written with the high T assemblage to the right ofthe equal sign

15) clin ta = H O sp anth 2

16) clin anth fa = herc en H O 2

9) fa ta = fs anth H O 2

10) clin fa ta = anth herc H O 2

12) clin anth = sp en H O 2

CFMASH

30) ta sp = H2O mgts anth

P(bar)

T CO

[atg chr fs di herc hed trfa H2O fta br mgts]

4

anth

en

ta

liz =

chr

(Che

rnos

ky, 1

975)

3.5 th

icken

ed pr

ism

5.5

thick

ened

pris

m

mk/C

01tneidarglamreht

nredoM

Ther

mal

gra

dien

t 25

C/k

m

Ancie

nt th

erm

al g

radi

ent

30 C

/Km

Fig. 8. A simplified P–T grid of Fig. 7 shows only the interesting metamorphic reactions. Maximum pressure during the cooling path is the pressure equivalent to the invariantpoint [atg chr fs di hed fa tr ta herc fta br mgts]. The vertical bar below the op. cit. invariant point shows the pressure equivalents to the Arabian–Nubian Shield’s ophioliticcrustal thicknesses as reconstructed by Stern et al. (2004). The two gray arrows show the modern thermal gradient and ancient thermal gradient. The black arrow showingthe path of the study samples, the peak pressure is 5–5.7 kbar based on the pressure calculation from Wadi Haimur-Abu Swayel ophiolites (Abd El-Naby and Frisch, 1999).Note: the ancient thermal gradient is equivalent to thickening of the sequence by factor of 5.5 as suggest by Ueno et al. (2011). The temperature axis is in two different scalesto show the reactions at high temperature condition in more details than Fig. 7.

12 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

indicators. Two water absent invariant points (508 �C to 1.08 kbarand 542 �C to 2.2 kbar) involve reactions with notable change inthe volume and can be used as pressure indicators. For better read-ing to the P–T grid, only the interesting reactions are shown inFig. 8 using two different scales for temperature axe.

6.2.1. Anthophyllite and talc formationOne of the key petrographic features is the relation between

pyroxene, anthophyllite and talc. The anthophyllite is a commonreplacement mineral of orthopyroxene. The anthophyllite can beformed due to eight metamorphic reactions (Fig. 8), however theabsence of clinochlore and the formation of the talc psuedomor-phic after anthophyllite make the only possibility to crystallizeanthophyllite is due to breakdown of high grade minerals (i.e.pyroxene). Two reactions can produce anthophyllite during a ret-rograde path at relatively high pressure (>1.7 kbar) and abovethe atg–chr–fs–di–hed–fa–tr–ta–herc–fta–br–mgts invariantpoint, however, these reactions produce clinochlore in consider-able values. This makes reaction fa + anth = fs + en + H2O and thelower pressure part (<1.7 kbar) of reaction fo + anth = en + H2Oare preferred way to produce anthophyllite in the assumed fluidcomposition.

Eight reactions can produce talc as a retrograde phase due tobreakdown of high grade assemblage that includes anthophyllite.Four reactions can be excluded since they contain clinochlore as a

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reactant or a product. The petrographic observation ‘‘orthopyrox-ene consumed due to talc growing’’ makes fa + ta = fs + anth + H2O,ta + sp = H2O mgts anth reactions (Fig. 8) are the favorable equilib-ria to produce talc. The two talc producing reactions have a temper-ature range 630–790 �C in a wide pressure condition. The pressureconditions of anthophyllite formation (<1.7 kbar) make the uppertemperature limit of talc producing reactions is below 730 �C. Otherreactions can produce anthophyllite and talc in the same pressure–temperature range but with different fluid compositions, thesereactions will be discussed in the fluid composition section.

Talc and anthophyllite formations indicate isobaric cooling pathat pressure below 1.7 kbar and in a temperature range of 550–800 �C. The cooling path can be extended to a lower temperaturecondition based on the presence of lizardite in the studied assem-blage. This conclusion is in agreement with the greenschist faciesconditions of the intermediate zone of the spinel grains (Fig. 4d).

Stern et al. (2004) reconstructed the ophiolitic sequence of theArabian–Nubian Shield and concluded that the ophiolitic succes-sions have crustal thicknesses of 2.5–5 km. These crustal thick-nesses are equivalent to pressure 0.7 and 1.4 kbar, respectively(Fig. 8) assuming lithostatic conditions and a rock density of2.84 � 103 kg/m3 (Carlson and Raskin, 1984). This constrains pres-sure conditions of the formation of the anthophyllite and talc pro-cess by 0.7–1.4 kbar (the retrograde path as shown by the blackarrow in Fig. 8).

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

Page 13

T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 13

6.2.2. Chrysotile formation and prograde metamorphismPresence of chrysotile fibers traversing the lizardite matrix indi-

cates that the rocks passed the reaction liz = chr (Fig. 8). Hamdy andLebda (2007) showed that the magnetite rims of the chromite grainsof Malo Grim serpentinites (part of the Sol Hamed ophiolites) equil-ibrated at a temperature range of 500–550 �C. These conditions arein agreement with the composition of the rim zones of the spinelgrains which show condition of amphibolite facies (Fig. 4d). Neitherpetrographic observations nor mineral chemistry data allow pre-dicting the pressure conditions of chrysotile formation.

The Arabian–Nubian Shield ophiolites were obducted withinvolcanic arc assemblages due to arc–arc collision process (e.g.Stern, 1994; Kusky et al., 2003; Meert, 2003; Stern et al., 2004).Obducted ophiolites, associated volcanics and sediments may rep-resent an accretionary prism system. Here we will follow theassumption of Valli et al. (2004) that average thermal gradient ofancient and modern accretionary prisms can be in the range of30 �C/km and 10 �C/km, respectively (Fig. 8). Abd El-Naby andFrisch (1999) studied Allaqi-Heiani ophiolite belt and they con-cluded that these ophiolites record temperature of 700 �C andpressures up to 8 kbar. These conditions can be converted to a ther-mal gradient of 25 �C/km which locates between the two assumedthermal gradient. This thermal gradient cuts the predicted temper-ature (500–550 �C) in a pressure range of 5.5–6.5 kbar (Fig. 8).

6.2.3. Fluid composition and T-XCO2 sectionDue to the ambiguity around the pressure condition during the

prograde path of the studied samples, the fluid composition will bestudied only along the cooling path.

Reaction equations are written with the high the equal sign

CFMASH-CO2

T CO

0.88100

150

200

250

300

350

400

450

500

550

600

650

700

750

0.90 0.92 0.9

14) ta sp = mgts anth H2O

18) ta e

15) sp anth = mgts en H2O

16) ta h17) her

12) ta fa = fs anth H2O 13) anth fa = en fs H2O

fs CO2 H2O

fta sid

fs sid fa CO2

1) ta sid = en fs CO2 H2O 2) fa ta = en fs H2O3) ta sid = en fa CO2 H2O4) ta mag = en CO2 H2O5) ta sid = fa mag CO2 H2O

5

6) en sid = fa mag CO2

[atg chr fsdi hed fasp herc

mgts fta sid]

7) ta mag = fo CO2 H2O

7

8) en mag = fo CO2

9) fo ta = en H2O

9

10) ta herc fta = fs mgts H2O

10

11) hed ta herc = fs di mgts H2O

1213

14

1819

20

19) ta f20) fo a

[atg chr fsdi hed fa

sp herc mgtsfta magsid CO ]2

21) sp 22) her

Fig. 9. A T-XCO2 grid in the system CFMASH-CO2 for the following end-members: anth,constructed at 1 kbar. Fluid concentration is buffered by the metamorphic reactions. Th

Please cite this article in press as: Abu-Alam, T.S., Hamdy, M.M. ThermodynaImplication for fluid interaction in the Arabian–Nubian Shield ophiolites. J. Afr

Fig. 9 shows a T-XCO2 grid in the system CFMASH-CO2 for thefollowing end-members: anth, atg, chr, en, fs, di, hed, fo, fa, ta,sp, herc, mgts, fta, mag, sid, H2O, CO2 at 1 kbar (the cooling pathof Fig. 8). The T-XCO2 grid was constructed in the full XCO2 range(not shown here), however all anthophyllite and talc producinginvariant points occur at high XCO2 (>0.88). In this type of section-ing (P-, T-XCO2), mineral phases are produced mainly at the invari-ant point conditions (Spear, 1993). The grid includes twenty-fiveunivariant reactions and seven invariant points. All of these invari-ant points occur at temperature range of 450–520 �C (Fig. 9). Allthe invariant points above 500 �C are magnesite–siderite absentinvariant points. At 500 �C and XCO2 (0.913), magnesite-bearinginvariant point appears. With cooling, the carbonate phase (sider-ite) becomes more stable (at 460 �C and XCO2 (0.978)). Below450 �C, the magnesite becomes metastable (Fig. 9). These invariantpoints show sequence of fluid evolution in the Sol Hamedserpentinites.

At XCO2 range (0.88–0.99), the first talc producing reaction(ta + sp = mgts + anth + H2O (Fig. 9)) is at higher temperature thanany anthophyllite producing reactions which were discussed in theP–T grid. Consequently reaction (herc + anth = mgts + en + fs + H2

O) is the preferred anthophyllite producing reaction. Once therocks started the cooling path, the anthophyllite producing reac-tion (op. cit.) buffers the fluid composition of the system and theT-XCO2 path (dashed arrows in Fig. 9) followed the reaction tillthe mineral composition arrives the atg–chr–di–hed–fo–fa–sp–fta–mag–sid–CO2 invariant point (510 �C; 0.998 (XCO2)). Theassemblage stayed at the invariant point conditions until one ofthe phases (i.e. fs, herc, mgts) was completely consumed or

T assemblage to the right of

4 0.96 0.98

n = anth H2O

erc = mgts anth fs H2O c anth = mgts en fs H2O

fs di CO2 H2Ohed ta sid

[atg chr dihed fo spherc mgtsfta mag]

1

2

3[atg chr fs dihed fo sp hercmgts fta]

4

6

8

11

15

16

17

o = anth H2Onth = en H2O

[atg chr hed fosp mgts fta

mag sid CO ]2

2122

ta = en mgts H2Oc ta = en mgts fs H2O

[atg chr fs di hedfo fa herc ftamag sid CO ]2

[atg chr dihed fo fa

sp fta magsid CO ]2

at 1 kbar

XCO2

atg, chr, en, fs, di, hed, fo, fa, ta, sp, herc, mgts, fta, mag, sid, H2O, CO2. The grid wase grid shows high CO2 concentration in the fluid.

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

Page 14

14 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx

excluded outside the equilibrium. At this stage of the path, therocks follow the isothermal reaction (ta + en = anth + H2O) whichproduces a considerable amount of talc. This reaction crosses allthe invariant points at 510 �C with different XCO2 composition(Fig. 9). Presence of magnesite in the studied assemblage (Table 4)and presence of magnesite-bearing invariant point at 500 �C andXCO2 (0.913) make the only possibility to terminate the talc pro-ducing reaction (op. cit.) is at the atg–chr–fs–di–hed–fa–sp–herc–mgts–fta–mag–sid–CO2 invariant point (510 �C; 0.885(XCO2)). The assemblage stayed at this invariant point until theanthophyllite was trapped and excluded outside the equilibrium,afterward the mineral equilibrium follows the reaction(fo + ta = en + H2O) until the magnesite-bearing invariant point at500 �C and XCO2 (0.913) which allows the first appearance of car-bonate-bearing phase. Forsterite consuming drives the equilibriumto leave the magnesite-bearing invariant point toward the magne-site–siderite-bearing invariant point (460 �C and XCO2 (0.978)).Subsequently the reaction (ta + sid = en + fa + CO2 + H2O) buffersthe equilibrium until the magnesite becomes metastable at450 �C and 0.984 (XCO2). Finally, reaction (ta + sid = en+ fs + CO2 + H2O) produces talc and siderite with constant consum-ing rate of H2O and CO2.

6.3. Fluid source and tectonic implications

Decarbonation of altered metabasalts and carbonates of marinesediments at low pressure condition has been considered as a pos-sible mechanism in order to explain CO2 fluxes at convergent

Ophioliteswith forearcsetting

Volcanic arc

Forearc basin

(a)

(b)

Accretionarprism

West-Gondwana

Mozambique Oc

CO

rich

flui

d2Mantle slices

Fig. 10. A three dimensional model illustrating the tectonic evolution of the studied ophfrom carbonate rocks in the subduction zone. These fluids re-concentrated in the forearcauthors recorded volcanic and sedimentary rocks in the Arabian–Nubian Shield were formThrusting and duplex thickening of the ophiolitic sequence. The white star is the positio

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margins (Staudigel et al., 1996; Kerrick and Connolly, 1998;Fischer et al., 1998; Molina and Poli, 2000). When hot geothermsare assumed, CO2-rich fluids can be transferred from the alteredoceanic crust to the mantle rocks in the subducted lithosphere(Fig. 10) in the forearc region (Molina and Poli, 2000). This mecha-nism can account for the CO2 enrichment of lithospheric mantle ona long-term scale and it may explain the occurrence of carbonatesin peridotite xenoliths (Ionov et al., 1993) as well as in some camp-tonitic lamprophyres (Bea et al., 1999). Here this mechanism canbe used to explain the high CO2 fluxes in the studied ophiolites(XCO2 = 0.89–0.99 (Fig. 9)). This high CO2 fluid content reactedwith the ophiolitic rocks in the forearc (Fig. 10) under pressurecondition of 1 kbar and temperature of around 800 �C (Fig. 8).Stern and Gwinn (1990) argued on the basis of C and Sr isotopesthat carbonate intrusions in the Eastern Desert of Egypt – whichcould be related to the carbonatizing fluids affecting Arabian–Nubian Shield ultramafic rocks – are mixtures of mantle derivedand remobilized sedimentary carbonate. Hamdy and Lebda(2007) gave the same conclusion based on carbon isotope compo-sition of south Eastern Desert of Egypt.

T-XCO2 grid (Fig. 9) shows that the fluid composition was buf-fered all the time by the metamorphic reactions (e.g. Greenwood,1975; Rice and Ferry, 1982; Spear, 1993; Abu-Alam et al., 2010).Field, petrographical and mineral chemistry evidences support thisthermodynamic observation. Majority of the T-XCO2 path tookplace at a temperature range of 450–550 �C. Most of the reactionsin this range of the temperature occurred as isothermal reactions,which means that the rocks were held at this temperature for a

Ophiolites

with mid-oceanic

ridge setting

passive margin

yOceanic crust

East-Gondwana

East-Gondwana

ean

iolites. (a) Development of subduction zone. High concentrated CO2 fluid is releasedophiolites. A passive margin is drawn on the flank of the oceanic basin since someed in a passive margin setting (e.g. Nakasib suture; Abdelsalam and Stern, 1993). (b)n of the studied ophiolites.

mic modelling of Sol Hamed serpentinite, South Eastern Desert of Egypt:. Earth Sci. (2014), http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001

Page 15

T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences xxx (2014) xxx–xxx 15

time period enough to consume one phase or more to drive theequilibria toward lower temperature conditions. Fig. 5a ofHamdy and Lebda (2007) shows that spinel minerals of the studiedophiolites were re-equilibrated at temperature condition of 500–550 �C which is the same range provided by the T-XCO2 grid. Pres-ence of magnesite in considerable amount in thin-section scale aswell as presence of small pockets and veins of magnesite in out-crop scale, indicate that the rocks were held for a long time atthe two magnesite-bearing invariant points (at temperature 500and 460 �C (Fig. 9)).

The high pressure condition (8 kbar) which was assumed byAbd El-Naby and Frisch (1999) and which was used here to predictthe geothermal gradient and the prograde path (the black arrow ofFig. 8) as well as the predicted pressure range (5.5–6.5 kbar fromthis study) can be explained in the context of extensive duplexarray and thickness of the original ophiolitic sequence (e.g.Hirono and Ogawa, 1998; Ueno et al., 2011) that associated withsubduction process. Oceanic crust in a forearc setting can be over-loaded by obduction of a crust that formed in a mid-oceanic ridgeand the thrusting in the forearc crust itself (e.g. Kromberg type-Section, Barberton Greenstone Belt, South Africa; Grosch et al.,2012). Original thickness of the Arabian–Nubian Shield’s ophioliticsequence is 2.5–5 km (Stern et al., 2004). Following oceanic crustdensity of 2.84 � 103 kg/m3 (Carlson and Raskin, 1984), the studiedophiolites were overloaded by 20–28 km thickness of obductedand thrusted oceanic crust from both mid-oceanic and forearc set-tings. This is in agreement with thickness increase of the originalsequence by a factor in range of 5.6 and 11.2. The same thickeningfactors were suggested numerically by Ueno et al. (2011). Theophiolite within the Allaqi area appears to have been emplacedSSW-ward above a NNE dipping subduction zone (Kusky andRamadan, 2002; Abdelsalam et al., 2003; Abdeen andAbdelghaffar, 2011) soon after its formation, i.e., 730–697 Ma (Aliet al., 2010).

One of the open questions around the ophiolites of the Arabian–Nubian Shield is ‘‘when did the alteration take place? Is it before orafter the obduction? (Stern et al., 2004)’’. Clearly, petrographicobservations and thermodynamic modelling that are presentedhere give an answer to this question. The studied ophiolites showtwo segments of the P–T path; one is the isobaric cooling path atpressure condition of 1 kbar and the second is prograde path froma pressure 1 kbar up to 5.5–6.5 kbar (black arrow of Fig. 8). The iso-baric cooling path occurred under oceanic crustal thickness of3.5 km which means that the first stage of alteration took placebefore the obduction while the second stage occurred duringthrusting and the obduction processes (prograde metamorphism).In the present situation, the ophiolites are thrusted over volcanicarc-assemblage. The volcanic arc-assemblage of the Arabian–Nubian Shield records a peak pressure around 3–4 kbar (e.g.Noweir et al., 2006; Abu-Alam, 2005; Abu-Alam and Farahat,unpublished data). This can be ensued only if the ophiolitesachieved the peak condition (5.5–6.5 kbar) before the final thrust-ing above the low-pressure arc-assemblage.

7. Conclusions

The Sol Hamed serpentinised ophiolitic mantle peridotites inthe south Eastern Desert of Egypt at the Allaqi-Heiani-Onib-SolHamed-Yanbu arc–arc suture formed beneath the forearc in thesubducted slab. They are mainly harzburgites formed after highpartial melting of >20% in reducing conditions. These harzburgitesinteracted with Ti-rich melts (boninite) in SSZ, which latter formedthe Sol Hamed cumulates. The Sol Hamed peridotites later thrustedover low-grade arc-assemblage of the Arabian–Nubian Shield. Theyshow a P–T path of an isobaric cooling at lithostatic pressure of

Please cite this article in press as: Abu-Alam, T.S., Hamdy, M.M. ThermodynaImplication for fluid interaction in the Arabian–Nubian Shield ophiolites. J. Afr

1 kbar which is equivalent to an oceanic crustal thickness of3.5 km. The alteration occurred before the thrusting and at highCO2 fluxes. The decarbonation of altered oceanic metabasalts andcarbonates of marine sediments at low pressure condition can beconsidered as a possible mechanism to explain the high concen-trated CO2 fluid fluxes at the convergent margin. The concentrationof the fluid during the cooling path was buffered by the metamor-phic reactions. The second segment of the path represents a pro-grade metamorphism which occurred under extensive duplexarray and thrusting of the oceanic crust. The crust in the forearcbasin was overloaded by 20–28 km of obducted and thrusted oce-anic crust from both mid-oceanic and forearc basin. This is equiv-alent to thickness increase of the original ophiolitic sequence by afactor in range of 5.6 and 11.2.

Acknowledgements

All analyses were carried out via personal communications. Sin-cere thanks are due to Dr. Ryszard Orlowski and Mrs. Tatiana Wes-olowska (Institute of Geological Sciences-Polish Academy ofSciences) for help with the chemical analysis of minerals andwhole rock, respectively. We thank R. Stern and K. Stüwe for thediscussion around accretionary prisms. T. Holland is thanked forhis help to provide a new thermodynamic dataset includes lizar-dite. J. Connolly and F. Gallien are thanked for their help with Per-Ple-X program. B. Evans is thanked for his help with inaccessiblepapers. This paper has been greatly improved by R. Greiling, A.Fowler and the three anonymous reviewers.

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