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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Thermodynamic modelling of Sol Hamed serpentinite, South Eastern
Desert of Egypt: Implication for fluid interaction in the Arabian–Nubian
Shield ophiolites
Tamer S. Abu-Alam
a,b,c,
⇑
, Mohamed M. Hamdy
a
a
Geology Department, Faculty of Science, Tanta University, Tanta, Egypt
b
Norwegian Polar Institute, Hjalmar Johansens gt. 14, NO-9296 Tromsø, Norway
c
Egyptian Institute of Geodynamic, Cairo, Egypt
article info
Article history:
Received 20 March 2013
Received in revised form 1 June 2014
Accepted 3 June 2014
Available online 24 June 2014
Keywords:
Arabian–Nubian Shield
Forearc subducted peridotite
Ophiolites
Carbonatization
Thermodynamic modelling
T-XCO
2
abstract
The Arabian–Nubian Shield is the largest tract of juvenile continental crust of the Neoproterozoic. This
juvenile crust is composed of intra-oceanic island arc/back arc basin complexes and micro-continents
welded together along sutures as the Mozambique Ocean was closed. Some of these sutures are marked
by ophiolite decorated linear belts. The Sol Hamed ophiolite (808 ± 14 Ma) in southeastern Egypt at the
Allaqi-Heiani-Onib-Sol Hamed-Yanbu arc–arc suture represents an uncommon example of rocks that
might be less deformed than other ophiolites in the Arabian–Nubian Shield. In order to understand
fluid–rock interactions before and during arc–arc collision, petrological, mineral chemistry, whole-rock
chemistry and thermodynamic studies were applied to the Sol Hamed serpentinized ophiolitic mantle
fragment. These studies reveal that the protolith had a harzburgite composition that probably originated
as forearc mantle in the subducted oceanic slab. We propose that these rocks interacted with Ti-rich
melts (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 refractory
residues 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 two
stages of fluid–rock interaction. The first stage occurred as a result of infiltration of concentrated CO
2
-rich
fluid released from carbonate-bearing sediments and altered basalt at the subduction zone. The alteration
occurred during isobaric cooling at a pressure of 1 kbar. The fluid composition during the isobaric cooling
was buffered by the metamorphic reactions. The second stage of fluid–rock interactions took place
through prograde metamorphism. The increase in pressure during this stage occurred as a result of
thrusting within the oceanic crust. In this process the forearc crust was loaded by roughly 20–30 km
of overthrust rocks.
Ó2014 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 within
and around the margins of the Mozambique Ocean, which formed
in association with the breakup of Rodinia 800–900 Ma (Stern,
1994; Hassan et al., 2014). These crustal fragments collided as
the Mozambique Ocean closed around 600 Ma (Meert, 2003).
Due to this collision processes a supercontinent variously referred
to 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 rocks
separating less deformed volcanic arc rocks (Fig. 1).
The major suture zones of the ANS can be classified into two
types, arc–arc and arc-continent sutures (Abdelsalam and Stern,
1996). The arc–arc sutures trend mostly NE–SW representing the
zones 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-Yanbu
and Nakasib-Bir Umq sutures are good examples of this type. Fol-
lowing arc–arc collision, the ANS collided with pre-Neoproterozoic
http://dx.doi.org/10.1016/j.jafrearsci.2014.06.001
1464-343X/Ó2014 Elsevier Ltd. All rights reserved.
⇑
Corresponding author at: Norwegian Polar Institute, Hjalmar Johansens gt. 14,
NO-9296 Tromsø, Norway.
E-mail address: tamer.abu-alam@npolar.no (T.S. Abu-Alam).
Journal of African Earth Sciences 99 (2014) 7–23
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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 best
examples are the Nabitah suture in the east and the Keraf suture in
the west. The final collision of East- and West-Gondwana deformed
the ANS along north trending shortening zones developed between
650 and 550 Ma (Abdelsalam and Stern, 1996; Stern et al., 2004;
Abu-Alam and Stüwe, 2009). The well-known example of these
zones is the Hamisana shear zone, which is characterized by
east–west crustal shortening fabrics, steep folds, and thrust faults.
The sutures, typically reactivated as transpressional/transcurrent
zones, are located across the shield (e.g., Johnson et al., 2004, and
references therein). Late deformation included the occurrence of
lateral escape tectonics along transtensional or transpressional
systems during the final stages of orogeny (e.g., Stern, 1994;
Johnson et al., 2004; Meyer et al., 2014).
Ophiolitic rocks are remarkably abundant in the ANS. They are
scattered across most of the shield, over a distance of 3000 km
from the farthest north (Gebel Ess) almost to the equator, and from
Rahib in the west to Gebel Uwayjah (45°E) in the east (Fig. 1). The
abundance of the ophiolites is a further indication that the Ara-
bian–Nubian Shield was produced by processes similar to those
of modern plate tectonics (Stern et al., 2004). The ophiolitic rocks
of Eastern Desert (ED) of Egypt (Fig. 1) are interpreted to be formed
in a suprasubduction zone (SSZ) (e.g. Ahmed et al., 2006; Azer and
Stern, 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; El
Bahariya and Arai, 2003; Farahat et al., 2004). However, some stud-
ies suggested that these rocks have been generated in mid-ocean
ridges (Ries et al., 1983; Zimmer et al., 1995; El Bahariya and
Arai, 2003; Farahat, 2010). The East- and West-Gondwana collision
led to obduction of the SSZ ophiolitic rocks over a continental
Saudi Arabia
Al Amar
Idsas
Hulayfah
Najd Faults
B.Umq
Al
Wask
Thurwah
Red
Onib
Wadi Halfa
Sinai
Eastern
Desert
Egypt
Bayuda
Desert
Sudan
Nakasis
Baraka
Sea
Bishah
Hamdah
Closure
Butana
Qala
En Nahl
Nuba
Mtns Ingessana
hills
Ethiopia
Somalia
Marda F
Tulu
Dimitri
Didessa S
Kurmuk
Akobo S
Akabo Adola
Possible shear
Nyangea S
Aswa S
Karasuk Moyale
Mutilo F
Baragoi
Sekerr
Mt
Elgon Kenya
30 40 50
0
5
10
15
20
25
Pan-African sequences
(volcano-sedimentary arc
assemblage, syn-,
post-tectonic intrusions,
gneisses and schists)
Faults and shear zones
Ophiolite Complexes
25
20
10
15
30 40
Fig. 2
Gebel Ess
Gebel Uwayjah
Sol
Hamd
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).
8T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

margin (Akaad and Abu El Ela, 2002; El-Gaby, 2005) of the West-
Gondwana (Abd El-Rahman et al., 2009). Subduction was active
while the process of ophiolitic overthrusting was operative along
thrust planes (Kröner et al., 1987; Stern, 1994). The ophiolites
are not all the same age and there are progressive changes in age
of ophiolites across the ANS (Berhe, 1990; Shackleton, 1994; Ali
et al., 2010). They have an isotopic Neoproterozoic age ranging
from 890 to 690 Ma, documenting a 200 Ma year period of oceanic
magmatism, and are caught up in 780 Ma to 680 Ma suture
zones 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 an
evolving, subduction related continental arc setting; Miller and
Dixon, 1992), other younger rocks including dikes, molasse-type
sedimentary rocks, high-K volcanic rocks and alkaline granites
were formed during the last stages of the Neoproterozoic develop-
ment of the ANS. These rocks are related to a late extensional stage
developed in the ANS (Blasband et al., 2000; Johnson and
Woldehaimanot, 2003; Ghoneim et al., 2014).
Origin of the ultramafic protolith in SSZ or in mid-ocean ridges
and 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 exists
around origin and composition of the fluid during the alteration
process. Some authors suggested that the alteration of the
ultramafic rocks – the dominant component of the ANS’s ophiolites
– occurred by interaction with hot fluid during seafloor weathering
(e.g. Lebda, 1995; Li and Lee, 2006). Other authors believe that the
alteration took place by infiltration of metamorphic and hydro-
thermal fluid along major tectonic fractures during or after rock
exhumation (e.g. Hyndman and Peacock, 2003; Hamdy, 2004;
Hamdy and Lebda, 2007; Abu-Alam and Stüwe, 2011) or in the
subduction zone as the ultramafic rocks were a component of the
forearc (Hamdy et al., 2013).
Sol Hamed ophiolite in southeastern Egypt and northeastern
Sudan (Fitches et al., 1983) at the Allaqi-Heiani-Onib-Sol Hamed-
Yanbu arc–arc suture (Abdelsalam and Stern, 1996; Abdelsalam
et al., 2003) differs from other ophiolites further north in the ED
of 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 the
ophiolitic rocks north of the Sol Hamed represent a far-travelled
ophiolitic 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 deformed
by the movement along faults that occurred after the closure of
the Mozambique Ocean. In this work, petrological relationships,
mineral chemistry, geochemistry and thermodynamic modelling
are described and applied to ultramafic rocks from the Sol Hamed
ophiolite (Fig. 2). The results help to more clearly define the nature
of the fluid–rock interactions process occurred in the intraoceanic
collision of the ANS. This is essential for understanding the genesis
Serpentinite
Metagabbro
Pillow lava
Arc
metavolcanics
Mélange
metasediments
Fault
Thrust
01 2 km
22 19
22 14
36 12
36 05
W. El Kwan
W. Ditt
G.Makarib Sample location
Magnesite site
Chromite site
N
Fig. 2. Simplified geological map of Sol Hamed area modified after Abu El Laban (2002).
T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23 9

of the ANS Neoproterozoic ultramafic fluid–rock interactions and
its relation to mineralizations.
2. Geological setting
Many of the ultramafic outcrops in the Arabian–Nubian Shield
are detached, scattered and isolated (Fig. 1) due to intrusion of
syn- 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 existed
between 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 and
are 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 and
southern sectors of the Eastern Desert of Egypt (Fig. 1), where they
occur as tectonized bodies and mélanges of pillowed metabasalt,
metagabbro, and variably altered ultramafic rocks (El Sharkawy
and El Bayoumi, 1979). The ultramafic rocks are mostly serpenti-
nized with relicts of fresh ultramafic protolith, but sometimes
include quartz carbonates (i.e. listwaenite), talc-carbonates, mag-
nesite veins and chromite pods.
The Sol Hamed ophiolite is a part of Allaqi-Heiani-Onib-Sol
Hamed-Yanbu arc–arc suture (Abdelsalam and Stern, 1996;
Abdelsalam et al., 2003). This arc–arc suture is considered – along
with the Ariab-Nakasib-Thurwah-Bir Umq suture farther south in
Arabia and Sudan (Johnson et al., 2004) to be one of the two longest
and most complete Neoproterozoic ophiolite-decorated sutures in
the 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–west
trending axes to produce upright antiforms and late-stage south-
east verging thrusts. Vergence of the ophiolite nappe was used to
infer a north dipping subduction zone along the line of a suture
which 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 of
serpentinites, chromite-bearing serpentinites and magnesite-bear-
ing serpentinites, forming the base of a dismembered ophiolitic
sequence. The Sol Hamed ophiolite sequence comprises also
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 km
long and about 0.4–1.8 km wide, elongated in NE–SW direction.
They are bounded and thrusted over arc metavolcanics from the
northwest. Some rock portions are extremely altered, especially
along 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 bordered
with a significant mélange metasediments belt (up to 1 km wide).
The mélange metasediments include in addition to the serpentinite
bodies metavolcanics, metagabbros, schists and rare marble. This
mélange belt separates the tectonized serpentinite bodies from
the 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
Table 1
Summary of mineral assemblages of the studied ultramafic rocks.
Primary minerals Olivine
Opx
Spinel
Chromite (the inner core)
First stage of alteration and metamorphism Lizardite
Antigorite
Anthophyllite
Magnesite
Magnetite
Chromite (inter. zone)
Sulfides
Second stage of alteration and metamorphism Lizardite
Chrysotile
Talc
Magnesite
Magnetite
Chromite (outer zone)
Chlorite
Talc
Magnetite and spinel
(a)
(b)
Zoned Chromite
Chrysotile
Olivine
~ 5 mm
s
s
mag
mag
anth
anth
anth
anth
anth
opx opx
opx
liz
liz
liz
liz liz
liz
20 µm
Fig. 3. (a) Semi-schematic drawing showing the ophiolitic ultramafic of Sol Hamed
area, lizardite and anthophyllite are metamorphosed after orthopyroxene. Chrys-
otile is metamorphosed after lizardite. Talc is after anthophyllite. Liz, opx, anth, mag
and s are lizardite, orthopyroxene, anthophyllite, magnesite and sulfides, respec-
tively. (b) Zoned spinel in Sol Hamed serpentinite. Darker zones are richer in Cr.
10 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

the metagabbro. It is noteworthy that mapping of the mélange belt
by Abu El Laban (2002) differs from the earlier maps of Fitches
et al. (1983) and Hussein (1981) as the contact between ultramafic
and 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 ultramafic
belt are found. These dunite bodies mainly contain chromitite
pods. In addition, the Sol Hamed serpentinites contain rare pyrox-
enite bodies that can be easily distinguished from its high content
of the pyroxene relics. Due to pervasive serpentinization, field
identification of the Sol Hamed ultramafics as mantle or cumulate
material was not possible – although the presence of chromite
masses in some serpentinite masses might suggest that at least
part of the ultramafic sequence is of cumulate origin (Church,
1981). Fitches et al. (1983) recorded that the Sol Hamed ophiolite
includes 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 mostly
derived from dunite. Micro-lenses and thin planar segregations
occur in the serpentinized peridotite. Gradual contacts between
massive 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, to
massive coarse-grained chromite ore. The highly deformed chro-
mite bodies are the most abundant. Magnesite veins cross-cut
the eastern periphery of the serpentinite rocks. Hamdy (2007),
based on the C–O isotopes of these veins, estimated that carbon
was supplied from both geothermal fluids (giving magnesite with
d
13
C values from –2.06‰to –4.34‰VPDB) and metamorphic car-
bonaceous sediments (giving magnesite with d
13
C values from –
9.44‰to –10‰VPDB).
3. Petrography
Variable degrees of alterations are observed in the studied
ultramafic 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 mainly
pseudomorphic mesh and bastite textures after olivine and ortho-
pyroxene and sometimes occurs as interlocking and penetrating
grains (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 of
fine-grained magnetite, which concentrates in veins cutting zoned
chromite (Fig. 3a) or along relict pyroxene cleavages. Pyroxene
relicts occur as inclusions in anthophyllite (Fig. 3a). The anthophyl-
lite is a common replacement mineral of orthopyroxene, where it
initially grows along cleavage planes and eventually replaces the
whole grain. Talc is not abundant in the studied serpentinites. It
forms fine shreds, dense fibers and medium grained flaky crystals
(0.01–0.04 mm). Perfect cleavage, straight extinction and high
interference colors are characteristic features of the talc. The talc
is 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 sulfide
grains. Chromite occurs as disseminated subhedral and anhedral
crystals of reddish brown color. Some chromite grains look homo-
geneous in reflected light. The carbonate minerals are optically
negative 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. magnesite
and siderite). The spinel minerals are opaque crystals with no
cleavage. The spinel crystals show zonation where the cores are
darker than the rims (Fig. 3b).
Table 2
Representative analyses of pyroxene and serpentine. The chemical formula of the serpentine minerals was calculated based on 28 oxygen atoms and ignoring the H
2
O, pyroxene
formula was calculated based on 6 oxygen atoms.
Sample Opx Serpentine
Liz Chr
325/1 325/2 118/2 118/3 118/4 369 333/1 333/2 347/1 347/2 278/1 310 278/2 369 325/1 325/2
SiO
2
58.21 58.37 53.22 54.27 52.38 41.94 40.62 43.78 41.3 41.56 43.04 43.96 44.54 43.23 42.6 42.47
TiO
2
0.12 0.09 b.d.l b.d.l 0.06 0.03 0.12 0.08 b.d.l 0.01 0.15 0.08 0.07 b.d.l 0.01 b.d.l
Al
2
O
3
0.74 0.67 0.04 0.04 1.22 0.28 0.3 0.06 0.26 0.14 0.14 0.14 0.28 0.21 b.d.l 0.13
Cr
2
O
3
0.42 0.36 b.d.l b.d.l b.d.l 0.36 0.22 b.d.l 0.16 0.09 0.34 0.12 b.d.l 0.15 b.d.l 0.02
FeO 4.85 5.09 1.10 1.12 1.10 6.21 4.2 2.66 3.36 3.99 1.85 0.94 1.37 2.11 2.05 1.81
MnO 0.03 0.07 0.01 0.01 b.d.l b.d.l 0.38 b.d.l 0.22 0.17 0.09 b.d.l 0.01 0.01 b.d.l b.d.l
MgO 34.99 35.1 18.92 19.29 20.22 35.98 34.37 37.74 38.68 38.06 36.88 38.14 37.78 37.17 36.68 38.42
CaO b.d.l 0.03 24.08 24.56 23.02 0.11 0.03 0.07 b.d.l b.d.l 0.13 b.d.l 0.06 0.18 0.19 b.d.l
Na
2
O b.d.l b.d.l 0.03 0.03 0.03 0.08 b.d.l 0.12 b.d.l b.d.l 0.01 0.1 0.01 0.09 b.d.l 0.02
K
2
O b.d.l b.d.l b.d.l b.d.l b.d.l 0.01 b.d.l 0.01 b.d.l 0.01 0.03 0.08 0.02 0.08 0.03 0.01
Total 99.36 99.78 97.39 99.32 98.02 85 80.24 84.52 83.98 84.03 82.66 83.56 84.14 83.23 81.56 82.88
Si 2 2 1.974 1.974 1.927 8.162 8.283 8.369 8.036 8.097 8.385 8.418 8.47 8.375 8.41 8.258
Ti 0.003 0.002 – – 0.002 0.004 0.018 0.012 – 0.001 0.022 0.012 0.01 – 0.001 –
Al 0.03 0.027 0.002 0.002 0.053 0.064 0.072 0.014 0.06 0.032 0.032 0.032 0.063 0.048 – 0.03
Cr 0.011 0.01 – – – 0.055 0.035 – 0.025 0.014 0.052 0.018 – 0.023 – 0.003
Fe 0.139 0.146 0.029 0.031 0.028 1.011 0.716 0.425 0.547 0.65 0.301 0.151 0.218 0.342 0.338 0.294
Mn 0.001 0.002 – – – – 0.066 – 0.036 0.028 0.015 – 0.002 0.002 – –
Mg 1.792 1.792 1.046 1.046 1.109 10.43 10.44 10.75 11.21 11.05 10.71 10.88 10.71 10.73 10.79 11.13
Ca – 0.001 0.957 0.957 0.908 0.023 0.007 0.014 – – 0.027 – 0.012 0.037 0.04 –
Na – – 0.002 0.002 0.002 0.03 – 0.044 – – 0.004 0.037 0.004 0.034 – 0.008
K – – – – – 0.002 – 0.002 – 0.002 0.007 0.02 0.005 0.02 0.008 0.002
Cations 3.977 3.98 4.01 4.01 4.03 19.79 19.65 19.64 19.92 19.88 19.56 19.58 19.49 19.62 19.59 19.73
di – – 0.94 0.94 0.92
hed – – 0.002 0.002 0.001
en 0.93 0.92 – – –
fs 0.0055 0.006 – – –
mgts 0.027 0.01 – – –
T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23 11

90
90
90
90
Mg
Fe+Mn Al+Cr
Mg
(a)
50
0
50
Fe
Cr
Granulite-facies
Spinels
Upper-amphibolite-
facies spinels
Lower-amphibolite-
facies spinels
Fe-spinels
facies Cr-spinels
Al
3+
Greenschist-
3+
3+
(d)
Core intermediate zone rim
ab
c
d
e
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# variation
diagram for the investigated spinel. field I represents Cr-spinels of mantle peridotite, field II represents magnetite from metamorphic rocks and field III is magnetite from
unmetamorphosed igneous rocks (fields after Roeder, 1994; Mondal et al., 2001), (c) MgO–CaO–FeO ternary diagram showing the chemical composition of the carbonate
minerals, (d) compositional changes in spinels expressed in a triangular Cr–Fe
3+
–Al
3+
plot with reference to the fields of spinel types: a – chromian magnetite, b – aluminian
magnetite, 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).
12 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

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 electron
microscope equipped with Link Analytical AN-1000/855 energy
dispersive X-ray spectrometer. The analytical conditions were
15 kV accelerating voltage and 35 nA beam current. The following
standards were used during the measurements: enstatite for MgO
and SiO
2
; adularia for K
2
O and Al
2
O
3
; ferrosilite for FeO; titanite for
TiO
2
and CaO; jadeite for Na
2
O; spessartine for MnO; chromite for
the Cr
2
O
3
. Mineral formula and activity of the end-members were
calculated by AX program (http://www.esc.cam.ac.uk/research/
research-groups/holland/ax). The mineral abbreviations which will
be used in the following sections are from Holland and Powell
(2011).
CaO content is below 0.03 wt% in the orthopyroxene and FeO
content is in the range of 4.85–5.09 wt% while MgO content is
around 35 wt% (Table 2). This reveals that the main orthopyroxene
end-member is enstatite. While diopside is the main end-member
for the clinopyrocene. SiO
2
content of the serpentine ranges
between 40.62 and 44.54 (Table 2). Al
2
O
3
is in the range of below
the detection limit up to 1.79 wt%. The serpentines are classified to
lizardite 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 between
34.37 to 38.68 wt%. The MgO and the FeO ranges indicate ionic
substitution between Fe
2+
and Mg
2+
(Fig. 4a). FeO and Cr
2
O
3
con-
tents in lizardite increase (0.94–6.21 wt% and from below detec-
tion limit to 0.36 wt% for FeO and Cr
2
O
3
, respectively) distinctly
with 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 therefore
remain in their original crystal lattice.
Low Al
2
O
3
and TiO
2
contents in talc chemistry reveal limitation
in substitution between Si, Ti and Al (Table 3). The activities of talc
and Fe-talc end-members are in the range of 0.68–0.85 and
0.00013–0.00061, respectively. Table 4 shows chemical analyses
of the carbonate minerals. The high concentrations of MgO and
FeO (35.53–40.14 and 8.46–14.1 wt%, respectively) indicate high
activity of the magnesite and the siderite end-members (Fig. 4c).
The CaO content is in the range of 0.04–0.27 wt% revealing low
activity of the calcite.
Three compositional zones are distinguished for the spinel min-
erals. Compositions of core, intermediate and rim zones are given
in Table 5 and plotted in Al–Cr–Fe
3+
triangle of Stevens (1944)
(Fig. 4d). Cores and intermediate zones have aluminian chromite
to ferritchromite composition. Composition of the outer rim is
Cr-magnetite which is nearly devoid of Al and lie along the Cr–
Fe
3+
sideline (Fig. 4d). All the core compositions have Cr# and
Mg# displayed in the mantle chromite field of Roeder (1994)
(Fig. 4b).
The variation in the spinel composition can be interpreted as a
result of chemical alteration under hydrothermal conditions
(Abzalov, 1998; Barnes, 2000; Proenza et al., 2004). The alteration
is accompanied by decrease in Al, Mg and Cr contents and conse-
quently increase in Fe
3+
and Fe
2+
. Apparently with the increasing
of the alteration, Fe releases from olivine and orthopyroxene and
Cr releases from chromite and are accommodated in the serpen-
tines. According to Barnes (2000), the Fe
3+
-rich aluminian chro-
mite–ferritchromite zone were formed by reactions between the
Table 3
Representative analyses of talc. The chemical formula and the end-members activities
were calculated by the AX program and based on 11 oxygen atoms and ignoring the
H
2
O.
Sample Talc
306/1 306/2 306/3 306/4 306/5
SiO
2
63.30 62.11 61.65 63.08 61.09
TiO
2
0.05 b.d.l 0.04 0.07 0.22
Al
2
O
3
0.03 0.14 0.28 0.05 b.d.l
Cr
2
O
3
0.02 0.14 0.26 0.03 b.d.l
FeO 3.43 3.73 3.90 2.85 3.11
MnO b.d.l b.d.l 0.04 b.d.l b.d.l
MgO 28.13 27.84 27.17 28.57 27.60
CaO 0.09 0.02 b.d.l 0.16 0.16
Na
2
O 0.13 b.d.l 0.16 0.18 0.13
Totals 95.20 94.00 93.50 95.00 92.31
Si 4.053 4.036 4.034 4.041 4.035
Ti 0.002 – 0.002 0.004 0.011
Al 0.002 0.011 0.022 0.004 –
Cr 0.001 0.007 0.014 0.002 –
Fe
3+
–––––
Fe
2+
0.184 0.203 0.213 0.153 0.172
Mn – – 0.002 – –
Mg 2.685 2.696 2.65 2.728 2.717
Ca 0.006 0.001 – 0.011 0.011
Na 0.017 – 0.021 0.022 0.017
Sum 6.951 6.955 6.957 6.964 6.962
ta 0.71 0.72 0.68 0.75 0.74
fta 0.00023 0.00031 0.00036 0.00013 0.00019
Table 4
Carbonate analyses. The chemical formula was calculated based on 2 cations.
Sample Carbonate minerals
300/1 300/2 300/3 306/1 306/2 306/3 306/4 306/5 306/6 368/1 368/2 368/3 368/4
SiO
2
0.19 0.15 0.04 0.11 0.23 0.18 0.05 0.04 0.09 0.13 0.16 0.1 0.09
Cr
2
O
3
0.11 0.04 0.09 b.d.l 0.07 b.d.l b.d.l b.d.l b.d.l 0.05 0.04 0.02 0.07
FeO 13.47 11.17 13.42 14.1 12.33 8.46 13.03 13.08 13.79 8.97 13.06 13.95 10.63
MnO 0.09 0.25 0.13 0.33 0.33 0.22 0.03 0.08 0.17 0.24 0.25 0.25 0.27
MgO 36.19 38.06 36.82 35.53 36.53 40.14 36.57 36.09 36.63 39.53 36.58 36.08 38.07
CaO 0.11 0.23 0.17 0.15 0.15 0.24 0.04 0.21 0.27 0.17 0.21 0.21 0.15
Totals 50.16 49.9 50.67 50.22 49.64 49.24 49.72 49.5 50.95 49.09 50.3 50.61 49.28
Si 0.006 0.004 0.001 0.003 0.007 0.005 0.002 0.001 0.003 0.004 0.005 0.003 0.003
Cr 0.003 0.001 0.002 – 0.002 – – – – 0.001 0.001 0 0.002
Fe
2+
0.343 0.28 0.338 0.361 0.315 0.21 0.333 0.336 0.346 0.224 0.331 0.354 0.269
Mn 0.002 0.006 0.003 0.009 0.009 0.006 0.001 0.002 0.004 0.006 0.006 0.006 0.007
Mg 1.643 1.701 1.65 1.622 1.663 1.772 1.664 1.653 1.638 1.759 1.65 1.63 1.715
Ca 0.004 0.007 0.005 0.005 0.005 0.008 0.001 0.007 0.009 0.005 0.007 0.007 0.005
Sum2222222222222
mag 0.84 0.86 0.84 0.83 0.84 0.89 0.85 0.84 0.83 0.89 0.84 0.83 0.87
sid 0.26 0.21 0.25 0.27 0.24 0.17 0.25 0.25 0.26 0.18 0.25 0.26 0.21
T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23 13

aluminian chromite grains and surrounding magnetite rims, where
its 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 Hamed
ophiolitic peridotites is similar to that determined for many meta-
morphosed ultramafic rocks in the Eastern Desert (e.g. Ghoneim
and Aly, 1986; Azer and Khalil, 2005; Hamdy and Lebda, 2007).
Ultramafic bodies in orogenic belts are commonly metamorphosed
to the same grade as the surrounding rocks (Winter, 2001) and are
described by Evans (2004) as ‘‘isofacial’’. The metamorphic facies
that prevailed during the evolution of the Egyptian Neoproterozoic
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 trace
elements were carried out at the geochemistry laboratory of the
IGS-PAS. Concentrations of major and trace elements were deter-
mined after microwave-assisted acid digestion with atomic
absorption spectrophotometer (AAS-PU 9100xUNICAM). Before
Table 5
Spinel 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
SiO
2
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 0.04
TiO
2
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 0.11
Al
2
O
3
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 6.49
FeO 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 30.7
Cr
2
O
3
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 56.84
MnO 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 0.25
MgO 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 5.17
NiO 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.l
Total 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 0.011
Al 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 2.17
Ti 0.08 0.049 0.045 0.038 – 0.047 0.058 – 0.033 – 0.006 0.045 0.028 – 0.017 0.023
Cr 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 12.744
Mn 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 0.06
Mg 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 2.188
Ni – 0.16 0.11 0.04 0.11 – 0.05 0.23 – – 0.07 – 0.06 0.12 0.08 –
Fe
2+
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 5.752
Fe
3+
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 1.052
Mg# 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 0.275
Cr# 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 0.798
Fe
3+
# 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 0.0018
cmt 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
Table 6
Representative whole-rock chemistry of Sol Hamed serpentinites. Major oxides are in wt%, trace elements are in ppm. b.d.l is below detection limit.
Sample 306 310 323 324 325 333 340 347 369 378
Major oxides and LOI (wt%)
TiO
2
0.06 0.06 0.01 0.04 0.02 0.02 0.02 0.02 0.01 0.03
SiO
2
41.49 41.49 39.14 38.60 39.76 39.54 39.30 38.41 38.55 41.26
Al
2
O
3
0.48 0.88 0.42 0.36 0.39 0.20 0.91 0.31 0.80 0.14
Fe
2
O
3
6.79 7.49 6.58 7.99 7.84 8.51 8.24 7.54 8.70 5.99
MnO 0.09 0.07 0.09 0.07 0.08 0.12 0.08 0.09 0.09 0.03
MgO 38.29 39.11 39.22 38.88 38.37 38.73 39.39 38.43 39.83 39.35
CaO 0.18 0.12 0.13 0.24 0.17 0.05 0.18 0.14 0.65 0.04
Na
2
O 0.09 0.17 0.03 0.01 0.01 0.00 0.22 b.d.l 0.01 0.01
K
2
O 0.01 0.05 0.01 0.00 0.01 0.01 0.03 b.d.l 0.01 b.d.l
Sum 87.49 85.94 85.61 86.15 86.63 87.17 88.37 86.00 88.65 86.82
LOI 10.60 13.14 13.49 13.20 13.07 11.27 11.46 12.99 10.27 12.81
Trace elements (ppm)
Cr 2742.20 2696.98 2706.00 2652.89 2677.80 2582.89 2699.83 2717.47 2703.71 2420.37
Co 166.79 120.41 119.56 162.00 164.55 150.42 156.28 154.23 116.20 135.18
Ni 2381.21 1799.16 2055.34 2055.90 2072.52 1836.13 1815.06 2060.51 1650.08 1996.13
Cu 63.83 14.97 23.89 51.52 42.15 53.09 54.68 8.27 46.32 32.97
Zn 57.11 13.32 25.68 23.09 33.25 18.15 37.80 16.55 11.98 23.06
Sr 55.04 46.61 48.12 61.96 59.05 72.02 49.89 89.11 48.08 47.93
V 40.31 29.83 27.41 37.11 25.98 22.15 39.77 14.00 33.93 19.66
Ba 35.24 15.00 24.98 15.14 19.88 17.06 16.34 45.00 20.00 29.88
Pb 4.70 13.15 15.75 14.96 17.32 19.88 17.00 4.63 24.04 14.40
Cd 2.00 2.00 3.31 3.30 2.13 2.11 3.31 3.31 3.43 2.88
Li 10.08 4.99 6.99 6.15 5.11 7.96 9.00 9.93 1.68 7.14
Rb 1.80 0.97 1.34 1.51 1.44 1.65 0.32 0.86 0.17 1.27
14 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

digestion samples were heated to 1100 °C to determine loss on
ignition (LOI). Analytical precision was better than 0.5% for major
elements and 4 ppm for trace elements.
Due to the almost complete serpentinization of some of the Sol
Hamed peridotites, modal compositions could not be determined.
Therefore, normative compositions were calculated from anhy-
drous analyses using the CIPW norm, assuming a Fe
2
O
3
/FeO ratio
of 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 aware
of the fact that serpentinization may severally change the original
silica, magnesium and calcium values. In addition, Al, Na and K
substitute Si, Fe, Mg and Ca in primary mantle silicates (especially
in 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 shift
from the lherzolite into the harzburgite. However, trace element
values 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-mobile
elements may be directly related to the serpentinization process
or due to metasomatism by subduction-related fluids (Hamdy
et al., 2013).
6. Discussion
6.1. Origin and tectonic setting of the serpentinite protolith
Earth contains two main shallow mantle domains: sub-oceanic
lithosphere and sub-continental lithosphere. The Sol Hamed
harzburgite falls within the oceanic array (Niu, 2004) in MgO/
SiO
2
–Al
2
O
3
/SiO
2
space (Fig. 5). The oceanic array is parallel to the
terrestrial array but offset to lower MgO/SiO
2
values, presumably
due to loss of MgO during low-temperature seafloor weathering
and not due to the serpentinization process itself (Snow and
Dick, 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 term
these suprasubduction zone (SSZ) peridotites (Pearce et al.,
1984); a group that incorporates peridotites from both island arcs
and spreading centers above subduction zones. These discrete
genetic 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 geotectonic
indicator (e.g. Barnes and Roeder, 2001). Chromium numbers [Cr/
(Cr + Fe
3+
+ Al)] higher than 0.6 are usually restricted to subduc-
tion-related rocks (Dick and Bullen, 1984). Ishii et al. (1992) used
the Mg# [Mg/(Mg + Fe
2+
)] and Cr# of the spinel to discriminate
between peridotites from MOR, forearc and back-arc settings. Spi-
nels from the Sol Hamed serpentinites lie in the chemical space of
the forearc peridotite (Fig. 6a) and distinctly higher than spinels
from MOR and back-arc basin in the Cr#. This indicates that the
Sol 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 setting
proposed by Church (1988) for the Sol Hamed ophiolite.
Despite this forearc nature of the Sol Hamed peridotites, TiO
2
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 TiO
2
contents at a given C#
and Al
2
O
3
content in spinels may be due to interaction with Ti-rich
melts (boninite) in SSZ. These melts might be those which formed
cumulates as a part of the ultramafic sequence at the Sol Hamed.
Melt-rock reaction produced boninitic melt and porous dunitic
channels 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 Hamed
chromitite deposit. According to the melt–rock interaction model
and despite the controversy concerning the importance of water
in the formation of chromitites, the Cr# of the spinel is controlled
by both the degree of depletion of the mantle source, due to previ-
ous melting, and the degree of the second-stage melting. The latter
is presumably controlled mainly by the melt/rock ratio, together
with temperature and compositions of the melt.
It is noteworthy that the Sol Hamed serpentinites are plotted
outside 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 southern
Eastern Desert of Egypt. The low Mg# at a given Cr# of spinel is
due to low equilibration temperature (Khedr and Arai, 2013). Com-
positions of the chromian spinel in the Sol Hamed serpentinites are
not similar with those of the chromian spinel in Arais serpentinites
of the south Eastern Desert of Egypt (Khedr and Arai, 2013). The
latter, in contrast, are similar to the compositions of the chromian
spinel in the serpentinites of the central Eastern Desert of Egypt.
However, Church (1988) based on the presence of wehrlite and
Iherzolite cumulates and the absence of troctolitic cumulates sug-
gested that the Sol Hamed ophiolite represents slices of primitive
suprasubduction zone that were slit by activity of strike-slip faults.
The association of forearc cumulate rocks that crystallized in
sequence of olivine–(chromite; high Cr)–clinopyroxene–orthopy-
roxene–plagioclase and boninitic volcanic rocks that crystallized
in sequence of olivine–orthopyroxene–clinopyroxene–plagioclase
with a ductile fault zone in the Troodos ophiolite (Murton, 1986)
and with a spreading center in the case of the Betts Cove ophiolite
of Newfoundland (Coish and Chrch, 1979; Coish et al., 1982) might
suggest that ophiolites with these characteristics owe their
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 infiltration
Sea-floor
weathering
Terrestrial melting array
Al2O3/SiO2
MgO/SiO2
Fig. 5. Whole rock MgO/SiO
2
–Al
2
O
3
/SiO
2
plot. The terrestrial array is a compilation
of subcontinental peridotites (Hart and Zindler, 1986) and represents a melt
depletion trend. The Sol Hamed serpentinites plot offset to lower MgO/SiO
2
values
because of alteration. Compositions of depleted MORB mantle (DMM; Workman
and Hart, 2005), primitive mantle (PM; McDonough and Sun, 1995) and seafloor
weathering trend (Snow and Dick, 1995) are plotted for comparison.
T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23 15

preservation in part to their origin as strike-slip fault slivers
detached from the frontal part of arcs as a result of oblique
subduction.
Yet, the forearc affinity of the studied serpentinized peridotites
imposes more debates on whether they were formed in mantle
wedge 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 between
forearc serpentinites from different settings. Based on REE patterns
and Ti content of these rocks, Deschamps et al. (2013) estimated
the nature of the initial protolith for serpentinites, as well as the
geological settings in which they were formed. The studied Sol
Hamed peridotites have Ti content analogous to that of the sub-
ducted slab serpentinites. Hellebrand et al. (2001) tested which
trace elements correlate with major element indicators of partial
melting in central Indian ridge peridotites. The most common of
these is the Cr# in spinel. They found a well-defined correlation
between moderately incompatible elements, such as heavy rare
earth elements (HREEs) in clinopyroxene with spinel Cr#.
Boninite
Back-arc
Mid-Ocean Ridge
100Mg/Mg+Fe
100Cr/Cr+Al
20
406080
20
40
60
80
Forearc
100
0
0
100
010203040 50
0.01
0.1
1
10
Al2O3 wt%
TiO2 wt%
MOR-type
peridotite
arc
high-Ti
Suprasubduction
zone peridotite
MORB
arc low-Ti
Cr#
TiO2 wt%
0 0.2 0.4 0.6 0.8 1.0
0
0.1
0.2
0.3
0.4
MOR
Forearc
peridotite
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 Dick
and 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 from
the 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; Farahat
et al., 2011). (b) TiO
2
vs. Al
2
O
3
(Kamenetsky et al., 2001). (c) Cr# vs. TiO
2
.(Barnes and Roeder, 2001). MORB, mid-ocean ridge basalt.
16 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

Hellebrand et al. (2001) developed an empirical equation (F=10ln
(Cr#) + 24) to estimate the degree of melting F(in percent) as a
function of spinel Cr#. Using the equation of Hellebrand et al.
(2001), the estimated melting in the studied peridotites ranges
from 20% to 22%. Pearce and Parkinson (1993) have shown that
modelling 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, and
HREE (incompatible). Using the whole-rock compositions of V
and MgO (anhydrous), the Sol Hamed peridotites are comparable
with those of Lee et al. (2003) and underwent melt extraction
26–29%. High melting degree of the Sol Hamed mantle reservoir
is in accordance with its formation in the forearc. Bonatti and
Michael (1989) proposed that mantle melting ranges from nearly
zero 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 recovered
from forearcs, which generally form during the early stages in
the 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 Sol
Hamed ophiolite reflect melt–wallrock interactions (Stern et al.,
2004).
Serpentinization plays a role in redox conditions of the mantle
which change the oxidation state of redox-sensitive elements. It
is recognized that, for a constant amount of Fe remaining, the
ratios of ((Fe
2
O
3
1000)/(Fe
2
O
3
+ FeO)) in bulk rock increase with
the degree of serpentinization. To determine redox conditions for
the protoliths of serpentinites formed after the Neoproterozoic
peridotites of study, we used the V–MgO composition proposed
by Lee et al. (2003), since we know that V records fO
2
during man-
tle melting (e.g. Lee et al., 2003; Pearce and Parkinson, 1993). The
Sol Hamed peridotites have compositions between QFM and QFM-
1 (QFM-1; this refers to log fO
2
(QFM) = log units relative to quartz–
fayalite–magnetite buffer). This estimates that melt extraction in
most peridotites was reducing conditions. This is in agreement
with the serpentinites from the subduction zone (Parkinson and
Arculus, 1999).
6.2. Thermodynamic modelling
All the thermodynamic calculations in the following sections
were calculated by THERMOCALC (Powell and Holland, 1988), Per-
Ple_X (Connolly, 1990) and using the internally consistent dataset
of Holland and Powell (2011). Lizardite bearing reactions which
were proved experimentally (i.e. liz = br + atg (Evans, 2004),
liz = chr (Chernosky, 1975), liz = ta + fo + clin + H
2
O(Caruso and
Chernosky, 1979)) will be only used (Fig. 7).
Fig. 7 shows a P–Tgrid in the system CFMASH for the following
end-members: atg, chr, en, fs, di, hed, fo, fa, anth, tr, clin, ta, sp,
herc, mgts, fta, br, H
2
O. Activity of the H
2
O is imposed to be the
unity therefore all the CO
2
bearing phases are not seen in this grid.
The P–Tgrid shows forty-six univariant equilibria, five invariant
points and three experimental lizardite bearing reactions. All the
H
2
O bearing univariant reactions show steep slope in the P–T
space. Consequently these reactions can be used as temperature
anth
6
8
sp fo anth
herc fo ta
anth H
2
O
H
2
O di tr fa
H
2
O herc tr fs
tr fo H
2
O
herc tr fa H
2
O
H
2
O tr fa
herc fa ta H2O
tr fs H
2
O
chr
hed atg ta
hed atg
hed fa ta
clin en fa
clin hed fs
di atg
en ta
fo ta
clin en
3
4
5
7
clin hed
fo ta H2O
atg
H2O herc di tr fa
herc fo ta
1
2
9
10
11
12
13
14
15
16
100 150 200 250 300 350 400 450 500 550 600 650 700 750
200
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 of
the 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
clin anth fa
4
17) di clin = sp tr fo H O 2
17
liz = br atg (Evan, 2004)
clin hed ta
liz = chr (Chernosky, 1975)
liz = ta fo clin H2O
(Caruso & Chernosky, 1979)
herc fo anth
clin en fa
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
fs clin
18
19
20
21
22
23
25
24
27
28
29
26
30
[atg chr fs di hed
fa tr clin sp herc
fta br mgts]
[atg chr fs di hed tr
ta sp fta br mgts]
[atg chr fs di hed fa
tr ta herc fta br mgts]
[atg chr fs di hed tr
sp H2O fta br mgts]
[atg chr fs di herc hed tr
fa H2O fta br mgts]
H
2
O fo atg
clin en
sp fo ta
P
(bar)
T C
O
Fig. 7. P–Tgrid in the system CFMASH for atg, chr, en, fs, di, hed, fo, fa, anth, tr, clin, ta, sp, herc, mgts, fta, br, H
2
O. Activity of the H
2
O is imposed to be the unity. Note:
reactions liz = br + atg, liz = chr, liz = ta + fo + clin + H
2
O are used here after Evans (2004), Chernosky (1975) and Caruso and Chernosky (1979), respectively. Mineral
abbreviations are after Holland and Powell (2011).
T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23 17

indicators. Two water absent invariant points (508 °C to 1.08 kbar
and 542 °C to 2.2 kbar) involve reactions with notable change in
the volume and can be used as pressure indicators. For better read-
ing to the P–Tgrid, only the interesting reactions are shown in
Fig. 8 using two different scales for temperature axe.
6.2.1. Anthophyllite and talc formation
One of the key petrographic features is the relation between
pyroxene, anthophyllite and talc. The anthophyllite is a common
replacement mineral of orthopyroxene. The anthophyllite can be
formed due to eight metamorphic reactions (Fig. 8), however the
absence of clinochlore and the formation of the talc psuedomor-
phic after anthophyllite make the only possibility to crystallize
anthophyllite 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 above
the atg–chr–fs–di–hed–fa–tr–ta–herc–fta–br–mgts invariant
point, however, these reactions produce clinochlore in consider-
able values. This makes reaction fa + anth = fs + en + H
2
O and the
lower pressure part (<1.7 kbar) of reaction fo + anth = en + H
2
O
are preferred way to produce anthophyllite in the assumed fluid
composition.
Eight reactions can produce talc as a retrograde phase due to
breakdown of high grade assemblage that includes anthophyllite.
Four reactions can be excluded since they contain clinochlore as a
reactant or a product. The petrographic observation ‘‘orthopyrox-
ene consumed due to talc growing’’ makes fa + ta = fs + anth + H
2
O,
ta + sp = H
2
O 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 pressure
conditions of anthophyllite formation (<1.7 kbar) make the upper
temperature limit of talc producing reactions is below 730 °C. Other
reactions can produce anthophyllite and talc in the same pressure–
temperature range but with different fluid compositions, these
reactions will be discussed in the fluid composition section.
Talc and anthophyllite formations indicate isobaric cooling path
at pressure below 1.7 kbar and in a temperature range of 550–
800 °C. The cooling path can be extended to a lower temperature
condition based on the presence of lizardite in the studied assem-
blage. This conclusion is in agreement with the greenschist facies
conditions of the intermediate zone of the spinel grains (Fig. 4d).
Stern et al. (2004) reconstructed the ophiolitic sequence of the
Arabian–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 of
2.84 10
3
kg/m
3
(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 black
arrow in Fig. 8).
anth
8
sp fo anth
herc fo ta
anth H
2
O
clin en fa
en ta
fo ta
clin en
herc fo ta
2
9
10
12
15
16
clin anth fa
4
liz = br atg (Evan, 2004)
liz = ta fo clin H
2
O
(Caruso & Chernosky, 1979)
herc fo anth
clin en fa
30
[atg chr fs di hed
fa tr clin sp herc
fta br mgts]
[atg chr fs di hed tr
ta sp fta br mgts]
[atg chr fs di hed fa
tr ta herc fta br mgts]
[atg chr fs di hed tr
sp 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 of
the 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 C
O
[atg chr fs di herc hed tr
fa H2O fta br mgts]
4
anth
en ta
liz = chr (Chernosky, 1975)
3.5 thickened prism
5.5 thickened prism
mk
/C
0
1
tn
eid
a
rg
lamr
e
ht
n
r
e
d
o
M
Thermal gradient 25 C/km
Ancient thermal gradient 30 C/Km
Fig. 8. A simplified P–Tgrid of Fig. 7 shows only the interesting metamorphic reactions. Maximum pressure during the cooling path is the pressure equivalent to the invariant
point [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 ophiolitic
crustal thicknesses as reconstructed by Stern et al. (2004). The two gray arrows show the modern thermal gradient and ancient thermal gradient. The black arrow showing
the 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 scales
to show the reactions at high temperature condition in more details than Fig. 7.
18 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

6.2.2. Chrysotile formation and prograde metamorphism
Presence of chrysotile fibers traversing the lizardite matrix indi-
cates that the rocks passed the reaction liz = chr (Fig. 8). Hamdy and
Lebda (2007) showed that the magnetite rims of the chromite grains
of Malo Grim serpentinites (part of the Sol Hamed ophiolites) equil-
ibrated at a temperature range of 500–550 °C. These conditions are
in agreement with the composition of the rim zones of the spinel
grains which show condition of amphibolite facies (Fig. 4d). Neither
petrographic observations nor mineral chemistry data allow pre-
dicting the pressure conditions of chrysotile formation.
The Arabian–Nubian Shield ophiolites were obducted within
volcanic 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 the
assumption of Valli et al. (2004) that average thermal gradient of
ancient and modern accretionary prisms can be in the range of
30 °C/km and 10 °C/km, respectively (Fig. 8). Abd El-Naby and
Frisch (1999) studied Allaqi-Heiani ophiolite belt and they con-
cluded that these ophiolites record temperature of 700 °C and
pressures up to 8 kbar. These conditions can be converted to a ther-
mal gradient of 25 °C/km which locates between the two assumed
thermal 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-XCO
2
section
Due to the ambiguity around the pressure condition during the
prograde path of the studied samples, the fluid composition will be
studied only along the cooling path.
Fig. 9 shows a T-XCO
2
grid in the system CFMASH-CO
2
for the
following end-members: anth, atg, chr, en, fs, di, hed, fo, fa, ta,
sp, herc, mgts, fta, mag, sid, H
2
O, CO
2
at 1 kbar (the cooling path
of Fig. 8). The T-XCO
2
grid was constructed in the full XCO
2
range
(not shown here), however all anthophyllite and talc producing
invariant points occur at high XCO
2
(>0.88). In this type of section-
ing (P-, T-XCO
2
), mineral phases are produced mainly at the invari-
ant point conditions (Spear, 1993). The grid includes twenty-five
univariant reactions and seven invariant points. All of these invari-
ant points occur at temperature range of 450–520 °C(Fig. 9). All
the invariant points above 500 °C are magnesite–siderite absent
invariant points. At 500 °C and XCO
2
(0.913), magnesite-bearing
invariant point appears. With cooling, the carbonate phase (sider-
ite) becomes more stable (at 460 °C and XCO
2
(0.978)). Below
450 °C, the magnesite becomes metastable (Fig. 9). These invariant
points show sequence of fluid evolution in the Sol Hamed
serpentinites.
At XCO
2
range (0.88–0.99), the first talc producing reaction
(ta + sp = mgts + anth + H
2
O(Fig. 9)) is at higher temperature than
any anthophyllite producing reactions which were discussed in the
P–Tgrid. Consequently reaction (herc + anth = mgts + en + fs + H
2
O) is the preferred anthophyllite producing reaction. Once the
rocks started the cooling path, the anthophyllite producing reac-
tion (op. cit.) buffers the fluid composition of the system and the
T-XCO
2
path (dashed arrows in Fig. 9) followed the reaction till
the mineral composition arrives the atg–chr–di–hed–fo–fa–sp–
fta–mag–sid–CO
2
invariant point (510 °C; 0.998 (XCO
2
)). The
assemblage stayed at the invariant point conditions until one of
the phases (i.e. fs, herc, mgts) was completely consumed or
Reaction equations are written with the high T assemblage to the right of
the equal sign
CFMASH-CO2
T C
O
0.88
100
150
200
250
300
350
400
450
500
550
600
650
700
750
0.90 0.92 0.94 0.96 0.98
14) ta sp = mgts anth H2O
18) ta en = anth H2O
15) sp anth = mgts en H2O
16) ta herc = mgts anth fs H2O
17) herc anth = mgts en fs H2O
12) ta fa = fs anth H2O
13) anth fa = en fs H2O
fs CO2 H2O
fta sid
fs di CO
2
H
2
O
hed ta sid
fs sid
fa CO2
[atg chr di
hed fo sp
herc mgts
fta mag]
1
1) ta sid = en fs CO2 H2O
2
3
2) fa ta = en fs H2O
3) ta sid = en fa CO2 H2O
[atg chr fs di
hed fo sp herc
mgts fta]
4) ta mag = en CO2 H2O
4
5) ta sid = fa mag CO2 H2O
5
6) en sid = fa mag CO2
6
[atg chr fs
di hed fa
sp herc
mgts fta sid]
7) ta mag = fo CO2 H2O
7
8) en mag = fo CO2
8
9) fo ta = en H2O
9
10) ta herc fta = fs mgts H2O
10
11) hed ta herc = fs di mgts H2O
11
12 13
14
15
16
17
18
19
20
19) ta fo = anth H2O
20) fo anth = en H2O
[atg chr fs
di hed fa
sp herc mgts
fta mag
sid CO ]
2
[atg chr hed fo
sp mgts fta
mag sid CO ]
2
21
22
21) sp ta = en mgts H2O
22) herc ta = en mgts fs H2O
[atg chr fs di hed
fo fa herc fta
mag sid CO ]
2
[atg chr di
hed fo fa
sp fta mag
sid CO ]
2
at 1 kbar
XCO2
Fig. 9. A T-XCO
2
grid in the system CFMASH-CO
2
for the following end-members: anth, atg, chr, en, fs, di, hed, fo, fa, ta, sp, herc, mgts, fta, mag, sid, H
2
O, CO
2
. The grid was
constructed at 1 kbar. Fluid concentration is buffered by the metamorphic reactions. The grid shows high CO
2
concentration in the fluid.
T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23 19

excluded outside the equilibrium. At this stage of the path, the
rocks follow the isothermal reaction (ta + en = anth + H
2
O) which
produces a considerable amount of talc. This reaction crosses all
the invariant points at 510 °C with different XCO
2
composition
(Fig. 9). Presence of magnesite in the studied assemblage (Table 4)
and presence of magnesite-bearing invariant point at 500 °C and
XCO
2
(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–CO
2
invariant point (510 °C; 0.885
(XCO
2
)). The assemblage stayed at this invariant point until the
anthophyllite was trapped and excluded outside the equilibrium,
afterward the mineral equilibrium follows the reaction
(fo + ta = en + H
2
O) until the magnesite-bearing invariant point at
500 °C and XCO
2
(0.913) which allows the first appearance of car-
bonate-bearing phase. Forsterite consuming drives the equilibrium
to leave the magnesite-bearing invariant point toward the magne-
site–siderite-bearing invariant point (460 °C and XCO
2
(0.978)).
Subsequently the reaction (ta + sid = en + fa + CO
2
+H
2
O) buffers
the equilibrium until the magnesite becomes metastable at
450 °C and 0.984 (XCO
2
). Finally, reaction (ta + sid = en
+fs+CO
2
+H
2
O) produces talc and siderite with constant consum-
ing rate of H
2
O and CO
2
.
6.3. Fluid source and tectonic implications
Decarbonation of altered metabasalts and carbonates of marine
sediments at low pressure condition has been considered as a pos-
sible mechanism in order to explain CO
2
fluxes at convergent
margins (Staudigel et al., 1996; Kerrick and Connolly, 1998;
Fischer et al., 1998; Molina and Poli, 2000). When hot geotherms
are assumed, CO
2
-rich fluids can be transferred from the altered
oceanic 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 CO
2
enrichment of lithospheric mantle on
a long-term scale and it may explain the occurrence of carbonates
in peridotite xenoliths (Ionov et al., 1993) as well as in some camp-
tonitic lamprophyres (Bea et al., 1999). Here this mechanism can
be used to explain the high CO
2
fluxes in the studied ophiolites
(XCO
2
= 0.89–0.99 (Fig. 9)). This high CO
2
fluid content reacted
with the ophiolitic rocks in the forearc (Fig. 10) under pressure
condition of 1 kbar and temperature of around 800 °C(Fig. 8).
Stern and Gwinn (1990) argued on the basis of C and Sr isotopes
that carbonate intrusions in the Eastern Desert of Egypt – which
could be related to the carbonatizing fluids affecting Arabian–
Nubian Shield ultramafic rocks – are mixtures of mantle derived
and remobilized sedimentary carbonate. Hamdy and Lebda
(2007) gave the same conclusion based on carbon isotope compo-
sition of south Eastern Desert of Egypt.
T-XCO
2
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 this
thermodynamic observation. Majority of the T-XCO
2
path took
place at a temperature range of 450–550 °C. Most of the reactions
in this range of the temperature occurred as isothermal reactions,
which means that the rocks were held at this temperature for a
Ophiolites
with forearc
setting
Ophiolites
with mid-oceanic
ridge setting
Volcanic arc
Forearc basin
(a)
(b)
passive margin
Accretionary
prism
Oceanic crust
West-Gondwana
East-Gondwana
East-Gondwana
Mozambique Ocean
CO rich fluid
2
Mantle slices
Fig. 10. A three dimensional model illustrating the tectonic evolution of the studied ophiolites. (a) Development of subduction zone. High concentrated CO
2
fluid is released
from carbonate rocks in the subduction zone. These fluids re-concentrated in the forearc ophiolites. A passive margin is drawn on the flank of the oceanic basin since some
authors recorded volcanic and sedimentary rocks in the Arabian–Nubian Shield were formed in a passive margin setting (e.g. Nakasib suture; Abdelsalam and Stern, 1993). (b)
Thrusting and duplex thickening of the ophiolitic sequence. The white star is the position of the studied ophiolites.
20 T.S. Abu-Alam, M.M. Hamdy / Journal of African Earth Sciences 99 (2014) 7–23

time period enough to consume one phase or more to drive the
equilibria toward lower temperature conditions. Fig. 5a of
Hamdy and Lebda (2007) shows that spinel minerals of the studied
ophiolites were re-equilibrated at temperature condition of 500–
550 °C which is the same range provided by the T-XCO
2
grid. Pres-
ence of magnesite in considerable amount in thin-section scale as
well as presence of small pockets and veins of magnesite in out-
crop scale, indicate that the rocks were held for a long time at
the two magnesite-bearing invariant points (at temperature 500
and 460 °C(Fig. 9)).
The high pressure condition (8 kbar) which was assumed by
Abd El-Naby and Frisch (1999) and which was used here to predict
the geothermal gradient and the prograde path (the black arrow of
Fig. 8) as well as the predicted pressure range (5.5–6.5 kbar from
this study) can be explained in the context of extensive duplex
array and thickness of the original ophiolitic sequence (e.g.
Hirono and Ogawa, 1998; Ueno et al., 2011) that associated with
subduction process. Oceanic crust in a forearc setting can be over-
loaded by obduction of a crust that formed in a mid-oceanic ridge
and 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 ophiolitic
sequence is 2.5–5 km (Stern et al., 2004). Following oceanic crust
density of 2.84 10
3
kg/m
3
(Carlson and Raskin, 1984), the studied
ophiolites were overloaded by 20–28 km thickness of obducted
and thrusted oceanic crust from both mid-oceanic and forearc set-
tings. This is in agreement with thickness increase of the original
sequence by a factor in range of 5.6 and 11.2. The same thickening
factors were suggested numerically by Ueno et al. (2011). The
ophiolite within the Allaqi area appears to have been emplaced
SSW-ward above a NNE dipping subduction zone (Kusky and
Ramadan, 2002; Abdelsalam et al., 2003; Abdeen and
Abdelghaffar, 2011) soon after its formation, i.e., 730–697 Ma (Ali
et 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 or
after the obduction? (Stern et al., 2004)’’. Clearly, petrographic
observations and thermodynamic modelling that are presented
here give an answer to this question. The studied ophiolites show
two segments of the P–Tpath; one is the isobaric cooling path at
pressure condition of 1 kbar and the second is prograde path from
a 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 of
3.5 km which means that the first stage of alteration took place
before the obduction while the second stage occurred during
thrusting and the obduction processes (prograde metamorphism).
In the present situation, the ophiolites are thrusted over volcanic
arc-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 ophiolites
achieved 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 in
the south Eastern Desert of Egypt at the Allaqi-Heiani-Onib-Sol
Hamed-Yanbu arc–arc suture formed beneath the forearc in the
subducted slab. They are mainly harzburgites formed after high
partial melting of >20% in reducing conditions. These harzburgites
interacted with Ti-rich melts (boninite) in SSZ, which latter formed
the Sol Hamed cumulates. The Sol Hamed peridotites later thrusted
over low-grade arc-assemblage of the Arabian–Nubian Shield. They
show a P–Tpath of an isobaric cooling at lithostatic pressure of
1 kbar which is equivalent to an oceanic crustal thickness of
3.5 km. The alteration occurred before the thrusting and at high
CO
2
fluxes. The decarbonation of altered oceanic metabasalts and
carbonates of marine sediments at low pressure condition can be
considered as a possible mechanism to explain the high concen-
trated CO
2
fluid fluxes at the convergent margin. The concentration
of 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 duplex
array and thrusting of the oceanic crust. The crust in the forearc
basin 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 a
factor 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 of
Sciences) for help with the chemical analysis of minerals and
whole rock, respectively. We thank R. Stern and K. Stüwe for the
discussion around accretionary prisms. T. Holland is thanked for
his 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 inaccessible
papers. This paper has been greatly improved by R. Greiling, A.
Fowler and the three anonymous reviewers.
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