ArticlePDF Available

Abstract and Figures

The Feiran–Solaf metamorphic complex of Sinai, Egypt, is one of the highest grade metamorphic complexes of a series of basement domes that crop out throughout the Arabian-Nubian Shield. In the Eastern Desert of Egypt these basement domes have been interpreted as metamorphic core complexes exhumed in extensional settings. For the Feiran–Solaf complex an interpretation of the exhumation mechanism is difficult to obtain with structural arguments as all of its margins are obliterated by post-tectonic granites. Here, metamorphic methods are used to investigate its tectonic history and show that the complex was characterized by a single metamorphic cycle experiencing peak metamorphism at ∼700–750 °C and 7–8 kbar and subsequent isothermal decompression to ∼4–5 kbar, followed by near isobaric cooling to 450 °C. Correlation of this metamorphic evolution with the deformation history shows that peak metamorphism occurred prior to the compressive deformation phase D2, while the compressive D2 and D3 deformation occurred during the near isothermal decompression phase of the P–T loop. We interpret the concurrence of decompression of the P–T path and compression by structural shortening as evidence for the Najd fault system exhuming the complex in an oblique transpressive regime. However, final exhumation from ∼15 km depth must have occurred due to an unrelated mechanism.
Content may be subject to copyright.
Exhumation during oblique transpression: The Feiran–Solaf
region, Egypt
Institut fu
¨r Erdwissenschaften, Universita
¨t Graz, Universita
¨tsplatz 2, A-8010 Graz, Austria (
ABSTRACT The Feiran–Solaf metamorphic complex of Sinai, Egypt, is one of the highest grade metamorphic
complexes of a series of basement domes that crop out throughout the Arabian-Nubian Shield. In the
Eastern Desert of Egypt these basement domes have been interpreted as metamorphic core complexes
exhumed in extensional settings. For the Feiran–Solaf complex an interpretation of the exhumation
mechanism is difficult to obtain with structural arguments as all of its margins are obliterated by post-
tectonic granites. Here, metamorphic methods are used to investigate its tectonic history and show that
the complex was characterized by a single metamorphic cycle experiencing peak metamorphism at 700–
750 C and 7–8 kbar and subsequent isothermal decompression to 4–5 kbar, followed by near isobaric
cooling to 450 C. Correlation of this metamorphic evolution with the deformation history shows that
peak metamorphism occurred prior to the compressive deformation phase D
, while the compressive D
and D
deformation occurred during the near isothermal decompression phase of the P–T loop. We
interpret the concurrence of decompression of the P–T path and compression by structural shortening as
evidence for the Najd fault system exhuming the complex in an oblique transpressive regime. However,
final exhumation from 15 km depth must have occurred due to an unrelated mechanism.
Key words: deformation phases; exhumation process; Pan-African Orogeny; P–T–D path; thermo-
dynamic modelling.
The Sinai Peninsula is the northern segment of the
Arabian-Nubian Shield, which comprises the north-
ern part of the East African Orogen. The East
African Orogen was assembled by accretion of an
intraoceanic island arc system associated with the
closing of the Mozambique Ocean during the Neo-
proterozoic (Engel et al., 1980; Gass, 1982; Kro
1984; Kro
¨ner et al., 1994; Fritz et al., 1996, 2002;
Katz et al., 2004). The entire arc assembly was then
thrust during the Pan-African orogeny over a cra-
tonic, gneissic basement, which is exposed in a series
of tectonic windows throughout northern Africa and
Arabian Peninsula (Fig. 1). This basement is often
considered to be of much older age than the arc
assembly (El-Gaby et al., 1990). A lack of knowledge
regarding the relationship between the low-grade arc
assembly and the much higher grade basement units
underneath severely limits our ability to reconstruct
the tectonic evolution of this region (Stern &
Manton, 1987).
One way to constrain this relationship is by studying
the metamorphic evolution of the gneissic basement
and correlating this with the structural events that
covered and later exhumed the gneisses from under-
neath the arc assembly. Such studies have been per-
formed in several of the gneissic windows in the
Eastern Desert of Egypt (e.g. Meatiq and Sibai core
complexes, Fritz et al., 2002) and also in Sinai (e.g. Kid
area: Abu El-Enen et al., 2003; Blasband et al., 1997,
2000; Brooijmans et al., 2003; Eliwa et al., 2008; Taba
area: Abu El-Enen et al., 2004; Eliwa et al., 2008).
These studies have shown that the basement domes
may be high grade equivalents of the low-grade arc
assembly, and that some of these domes are core
complexes in the classic sense. Several of these studies
have also shown that the exhumation of the gneiss
domes is associated with the activity of a major (more
than 1000 km long and hundreds of km wide) NW–SE
striking sinistral shear zone system called the Najd
fault system (Fig. 1). However, one of the highest
grade complexes in Egypt and the largest metamorphic
one in Sinai has not been studied metamorphically and
its exhumation history has not been explained: the
Feiran–Solaf metamorphic complex (FSMC) (Fig. 2).
This complex constitutes a narrow strip of gneisses in
the northwestern part of the exposed igneous and
metamorphic rocks in Sinai.
El-Gaby & Ahmed (1980) recognized that the
FSMC is made up of two asymmetric doubly-plunging
antiforms, namely the Feiran antiform and the Solaf
antiform, separated by synformal trough. This geo-
metry suggests that the FSMC may be a core complex
similar to those of the Eastern Desert of Egypt (e.g.
Fritz et al., 1996, 2002; Loizenbauer et al., 2001; Abd
El-Naby et al., 2008; Khudeir et al., 2008), but we will
show here that this is not the case.
J. metamorphic Geol., 2009, 27, 439–459 doi:10.1111/j.1525-1314.2009.00827.x
2009 Blackwell Publishing Ltd 439
This study investigates the metamorphic evolution
of Feiran–Solaf complex to constrain the tectonic
evolution of the area. A mineral equilibria approach is
used with petrogenetic pseudosections and relate the
derived metamorphic history to the detailed structural
map and structural evolution derived by El-Shafei &
Kusky (2003). Our tectonic model is then correlated
with independent field evidence, existing geochrono-
logical ages from the region and with other gneiss
domes elsewhere in the Arabian-Nubian Shield.
The FSMC is a NW trending elongate folded belt
40 km long and 5–11 km wide (Fig. 2) parallel to the
orientation of the Najd fault system. It is made up of
migmatitic biotite and hornblende gneisses, quartzo-
feldspathic gneisses and hornblende gneisses with
subordinate schists, amphibolite and calcsilicate rocks.
The complex was intruded and surrounded by a
number of granitic plutons, including pre-, syn- to
post-tectonic granitoids, diorite and large volumes of
dykes with mafic to felsic composition. These intru-
sions obliterate all structural relationships along the
margins of the complex. The northwestern part of the
complex is covered by Phanerozoic sedimentary rocks.
The complex can be divided into two zones: the Feiran
zone in the north-west and the Solaf zone in the south-
east. These two zones are separated by a diorite
intrusion (Fig. 2). The Feiran zone contains migmatitic
biotite and hornblende gneisses, hornblende gneisses,
quartzofeldspathic gneisses and locally some dolerites.
In contrast, the Solaf zone is mostly composed of
quartzofeldspathic gneisses (with very minor bands of
metagraywackes but with mappable portions of calc-
silicate rocks; Fig. 2). There are also less abundant
granitic gneisses and amphibolites as small lenses and
The rock units in the entire complex are systemati-
cally arranged around two doubly-plunging antiformal
structures: The Feiran antiform in the north-west and
the Solaf antiform in the south-east (Fig. 2). Stern &
Manton (1988) suggested that a thrust fault separates
the two antiforms, but El-Shafei & Kusky (2003) found
no evidence for such a fault. Migmatitic biotite and
hornblende gneisses occur in the core of the structure,
the hornblende gneisses structurally above and the
quartzofeldspathic gneisses above those. The calcsili-
cates generally appear to be located in the structurally
highest levels. This arrangement has been interpreted
as a stratigraphic succession by Ahmed (1981).
El-Gaby & Ahmed (1980) interpreted the Feiran
36 42
Red Sea
An Nakhil
Pan-African arc assembly
Gneiss complexes
Phanerozoic sequence
fault zone
Ar Rika
fault zone
Ad Damm
fault zone
Bir Umq-Nakasib
Fig. 1. Metamorphic gneiss complexes, Najd fault system and sutures of the northern part of the Arabian-Nubian Shield. Black
arrows show principal stress directions for the Najd fault system as a whole. Small white arrows show strain in individual gneiss
complexes. Note that this is extensional in the Eastern desert, but transpressive in the Feiran–Solaf metamorphic complex.
(Complied from various sources, including: Abdelsalam & Stern, 1996; Fritz et al., 1996; Johnson, 1998).
440 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
gneisses to represent a thick (>5000 m) sedimentary
succession with minor mafic intercalations. Several
authors have observed that metamorphism in the
FSMC appears to increase towards the core of the
structure, i.e. towards deeper structural levels. In
the migmatitic biotite and hornblende gneisses, the
metamorphic grade is reported to have reached upper
amphibolite facies conditions from 640 to 700 Cat
pressure from 4 to 5 kbar during the Pan-African
Orogeny (e.g. Eliwa et al., 2008). A zircon age derived
by Stern & Manton (1987) suggests 632 ± 3 Ma for
the high-grade migmatitic event. A K–Ar biotite
cooling age (Eliwa et al., 2008) suggests 617–594 Ma
as the time of cooling and exhumation.
Structure and kinematics
Most authors agree that the history of the Pan-African
deformation in Sinai can be described in terms of four
Proterozoic deformation phases D
(e.g. Eliwa
et al., 2008), where D
is the primary foliation forming
event interpreted to be related to an early Pan-African
extensional environment. D
and D
are shortening
phases forming the large scale geometry of the meta-
morphic complexes and D
is a gentle warping event.
For the FSMC, Fowler & Hassan (2008) suggested
that the major doubly-plunging antiforms building the
map scale geometry of the complex were formed
during D
(Fig. 2). As this is a somewhat different
interpretation from that of El-Shafei & Kusky (2003),
a summary of the structural evolution is presented
based on the literature and our own data below.
The earliest structure observed in the FSMC is the
primary lithological layering termed S
. Lithological
bands are from <1 cm to several metres in thickness and
are defined mainly by variation in the amounts and type
of mafic minerals. Good examples of this primary lith-
ological layering (S
) are thin continuous layers of
amphibolite (Fig. 3a) and calcsilicate rock within
quartzofeldspathic gneisses (Fowler & Hassan, 2008).
Fowler & Hassan (2008) as well as El-Shafei &
Kusky (2003) suggested that the dominant metamor-
phic foliation in the FSMC is S
While this nomen-
clature is retained here, we note that Abu El-Enen
Fig. 2. Location map (MC, metamorphic complex) and simplified geological map of the Feiran–Solaf metamorphic complex. Contacts
are modified from Ahmed (1981); El-Shafei & Kusky (2003) and Fowler & Hassan (2008). The white squares are the sample locations
(F1, F63, F93d & F39). Structural cross-sections, A–A¢through the Solaf zone and B–B¢are drawn sub-parallel to the main fold axes
of the Feiran and the Solaf antiforms. Lithological changes are often gradational and the mapped contacts are therefore approximate.
2009 Blackwell Publishing Ltd
F2 fold axe
F3axial plane
Metamorphic foliatio n
F2 axial plane
442 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
et al. (2003, 2004) suggested that the main foliation
forming event in the Taba and Kid belts of Sinai is D
Indeed, we found remnants of an earlier fabric as
inclusion trails in garnet in the FSMC. Nevertheless, as
these inclusion trails are rare, the main foliation event
is termed D
, keeping in mind that earlier phases may
have occurred.
formed a strong stretching lineation and
intrafolial F
-folds generally trending NW–SE. Shear
sense indicators indicate west and north-west directed
transport but Fowler & Hassan (2008) showed that the
principal strain during this event was vertical flattening
with stretching in both NW–SE and NE–SW direc-
tions. They also suggested that this extensional D
reflected the larger-scale extension related to the
breakup of Rodinia. In contrast, El-Shafei & Kusky
(2003) inferred that D
was responsible for the map
scale geometry of the FSMC. However, because the S
foliation is wrapped around the hinges of the map scale
antiforms, we follow here Fowler & Hassan (2008) and
suggest that the map scale geometry must be related to
later deformation phases.
The D
phase formed tight folds (F
) (Fig. 3b)
occasionally associated with a crenulation cleavage
). The F
-folds commonly refold F
-folds. These
structures are well preserved in the Solaf zone along
Wadi Solaf. The F
-folds trend parallel to the F
(NW–SE and NNW–SSE) and clearly formed due to
shortening in NE–SW direction. Because F
and F
axes are largely parallel, El-Shafei & Kusky (2003)
interpreted F
and F
to be both related to the same
deformation phase D
The third deformation phase D
formed inclined
open folds (F
) causing the map scale structure of the
complex and the Feiran and the Solaf antiforms
(Fowler & Hassan, 2008). El-Shafei & Kusky (2003)
interpreted this phase to be related to the Najd fault
system and associated it with NW–SE sinistral strike–
slip shear zones. These sinistral movements are well
recorded along the entire study area (Fig. 3c) and well
preserved in the calcsilicates. The F
hinges are roughly
parallel to the F
and the F
hinges trending shallowly
(NW–SE and NNW–SSE) parallel to the regional
trend of the map scale antiforms. The F
commonly refold F
-folds (Fig. 3b).
The D
deformation phase is an open warping event
that domed up the entire area and is associated with
NNW–trending quartz veins and gashes. F
-folds are
extremely rare and trend NE–SW, ENE–WSW and
E–W (Fig. 3d). The best example of an F
-fold is the
NE–SW synformal trough which separates the Feiran
and the Solaf antiforms (Fig. 2). The presence of
NNW-trending quartz veins and the gashes indicates
extension in NE–SW direction. It is possible that D
(shear zones along the margins of the complex) and D
are responsible for the exhumation of the region.
However, this is difficult to infer in the field as the
post- and syn-tectonic D
granitic bodies surround the
entire complex and obliterate all of its margins. Some
strong deformation features are developed along the
margins of the metamorphic complex, but these are
apparently related to the intrusion of the syn-tectonic
granites. For example, horizontal flattening shows
both sinistral and dextral movements and is especially
found at the end of Wadi Um-Takha and in the
calcsilicate rocks at the end of Wadi Dehest Abu-Talb
(Fig. 2 SE end of study region). During the Red Sea
and the Suez Gulf rifting, the entire study area was
deformed by brittle extensional structural features.
These features appear clearly as dyke swarms with
Phanerozoic age (e.g. Ahmed & Youssef, 1976;
El-Shafei & Kusky, 2003) and dextral strike–slip
faults. The best example of these faults is the fault
which cuts Gabal El-Banat and Gabal Serbal (post-
tectonic granites) (Fig. 2).
Most rocks of the FSMC are quartzofeldspathic
paragneisses with or without hornblende and with or
without migmatitic features. Aluminous metasedi-
ments are rare. Stern & Manton (1987) suggested that
the protoliths of the Feiran gneisses were probably
very immature, volcanogenic wackes or perhaps lithic
arenites inter-bedded with felsic tuffs.
The migmatitic biotite and hornblende gneisses
(henceforth called ÔmigmatitesÕ) (Fig. 3e) in the core of
Feiran zone are metatexites formed due to partial
melting (El-Shafei & Kusky, 2003). They are composed
of quartzofeldspathic leucosomes containing minor
biotite and hornblende and hornblende-rich meso-
somes containing minor biotite. In general, the
migmatites are pyroxene-free rocks, but rare pyroxene-
bearing migmatites occur. Complex structural patterns
are observed; include stromatic, folded stromatic,
ptygmatic and schlieren structures. The characteris-
tic feature of the migmatites is the abundance of
leucosomes that occur both layer parallel and cross-
cutting to the principal foliation (S
). The migma-
tites are intercalated with hornblende gneisses and
Fig. 3. Field photographs. (a) Two conformably inter-bedded amphibolite bands parallel to the sedimentary relict bedding of the
paragneisses in the Solaf zone. (b) Close up view showing F
isoclinal fold superimposed by F
open fold, looking NW. (c) Close
up view showing the D
sinistral strike–slip movement. (d) Two small-scale domes in the same orientation of the Feiran and the
Solaf doubly-plunging antiforms. The small-scale synform (F
), which separates the two domes, has the orientation of the NE–SW.
(e) The migmatites in the core of the Feiran antiform crop out in the more deeply eroded part along Wadi Feiran. (f) Close up view
showing the exfoliation appearance of the dolerite plugs at the northeastern part of the Feiran zone (Wadi Rummana). (g) Deformed
granitic dyke, the paragneisses around this dyke are rich in K-feldspar. (h) Syn-tectonic granites intruded perpendicular to the
metamorphic foliations (S
) at the end of Wadi Um-Takha.
2009 Blackwell Publishing Ltd
quartzofeldspathic gneisses. These gneisses lack mig-
matitic features. There is no contact between the mi-
gmatites and the granite intrusions surrounding the
FSMC suggesting that the migmatization process is
not related to these granites.
The most abundant rock units in the Feiran zone are
the hornblende gneiss and the quartzofeldspathic
gneiss (Fig. 2). Rare metagraywacke bands occur.
There is intercalation and repetition between these two
rock types, but the proportion of quartzofeldspathic
gneisses increases toward the outer-rim of the Feiran
antiform. In the Solaf zone, the quartzofeldspathic
gneisses are the predominant rock unit. The horn-
blende gneiss is fine- to medium-grained, strongly
foliated and dark in colour. The quartzofeldspathic
gneisses are brownish grey colour, fine- to medium-
grained and strongly foliated. The hornblende gneisses
consist of hornblende, plagioclase, minor orthoclase
and quartz with or without biotite. They also contain
subordinate amounts of chlorite, iron oxides, zircon,
apatite, titanite and epidote. The quartzofeldspathic
gneisses are free of hornblende but otherwise very
similar. In the hornblende and the quartzofeldspathic
gneisses, the presence of relict sedimentary bedding
(e.g. El-Gaby & Ahmed, 1980; El-Gaby et al., 1990;
El-Shafei & Kusky, 2003) suggests that they had a
sedimentary origin.
The calcsilicates running parallel to and in contact
with the granitoid rocks at the outer most rim of the
Solaf antiform (Fig. 2) are characterized by well-
developed garnet crystals up to several centimetres in
size. The matrix is fine-grained, massive and light to
dark-green in colour containing calcite, wollastonite,
diopside, Ca-amphibole, scapolite, plagioclase and
quartz with subordinate amounts of zoisite clinozoi-
Minor lithologies (with smaller occurrence than is
mappable on Fig. 2) include amphibolites, doleri te plugs
and granitic as well as mafic dykes. The amphibolites
occur as conformable inter-bedded bands, linear
bodies, and irregular lenses in the paragneisses of the
Solaf zone and in the syntectonic granitoids. They are
composed essentially of hornblende and plagioclase
with minor biotite and quartz. Titanite, apatite,
epidote and opaques are common accessory minerals.
El-Tokhi (1992) suggested an igneous origin for the
amphibolites, based on chemical evidence. But,
according to Fowler & Hassan (2008) and our field
evidence (Fig. 3a), the amphibolites are interpreted to
be of sedimentary origin. The dolerites occur as plugs
within the hornblende gneisses and the quartzofelds-
pathic gneisses at the very northeastern part of the
Feiran zone (Wadi Rummana) and consist of plagio-
clase, amphibole, pyroxene relics and iron oxides. The
amphibole and the pyroxene occur interstitially between
plagioclase crystals. Granitic dykes and pegmatitic veins
also cut the region. These are deformed and contain
inclusions from the surrounding rocks. The paragneisses
around these granitic dykes are enriched in K-feldspar.
The amount of K-feldspar in the country rocks
increasing toward these dykes (Fig. 3g).
The plutonic rocks in the study area are represented
by pre-, syn- and post-tectonic granites and diorite.
The pre- and the syn-tectonic granitoid rocks intruded
along the southern and eastern borders of the Solaf
zone. They extend further east forming numerous low-
relief hills characterized by intense shearing and they
contain many xenoliths of amphibolite, hornblende
gneisses and quartzofeldspathic gneisses (El-Shafei &
Kusky, 2003). In the eastern part of the study area,
these granites intruded along pre-existing oblique-slip
faults and parallel to the metamorphic foliation. In the
southern part of the study area, they intruded more or
less perpendicular to the metamorphic foliation
(Fig. 3h) and along the NE–SW normal faults. The
post-tectonic granites are characterized by the absence
of the dyke swarms and form large mountainous out-
crops bordering the complex along the eastern and the
southern parts of the Feiran zone (Gabal El-Banat and
Gabal Serbal). A large undeformed diorite body forms
an elongated NE-striking intrusion. El-Shafei & Kusky
(2003) suggested that this diorite intrusion may have
intruded along a pre-existing crustal weakness located
between the Solaf and Feiran zones. The entire area is
cut by a large number of dyke swarms ranging from
the Precambrian to the Cenozoic in age (El-Shafei &
Kusky, 2003).
The rock types of the Feiran–Solaf complex can be
grouped into four major types: (i) migmatitic biotite
and hornblende gneisses (migmatites); (ii) hornblende
gneisses; (iii) quartzofeldspathic gneisses and (iv)
calcsilicates. The quartzofeldspathic gneisses have a
very simple mineralogy (Fig. 4a) consisting of biotite,
quartz, plagioclase and Fe–Ti oxides and will not be
discussed further. The calcsilicates have a complicated
fluid history and are studied separately (Abu-Alam &
¨we, Unpublished data). Here, we concentrate on
the migmatites, the hornblende gneisses, and the minor
but nevertheless important metagraywacke inclusions
therein. The minerals were analysed at the Institute of
Earth Science, Karl-Franzens-Universita
¨t Graz, Aus-
tria, using a JEOL JSM-6310 scanning electron
microscope following standard procedures, operating in
EDS WDS mode at 5 nA beam current, accelerating
voltage 15 kV and duration time is 100 s. The chemical
formulae were calculated using the PET1.1 (Dachs,
2004). The mineral abbreviations which will be used in
the following sections are from Holland & Powell
Migmatitic biotite and hornblende gneisses (Migmatites)
The leucosomes are composed of quartz, plagioclase
and K-feldspar with or without biotite and the meso-
somes contain plagioclase, amphibole, quartz and
444 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
~ 750 °C
~ 5 kbar
old Mg-hb
~ 450-550 °C
young act
200 µm
Hornblende gneisses
Hornblende gneisses
bi, q, chl
(pre- ) S
1 foliation
Iron oxide
Quartzofeldspathic gneisses
Metamorphic foliation
Metagraywackes 1000 µm
1000 µm
1 mm 1 mm
0.02 µm1 mm
500 µm
Fig. 4. Photomicrographs of the main rock types. (a) The poor mineral assemblage of the quartzofeldspathic gneisses, (b) Clinopyroxene
of sample F39 is highly deformed and aligned along the S
foliation. (c) Hornblende porphyroblast of sample F39 has inclusions of
clinopyroxene. (d) Metamorphic foliation (parallel crystals of the hornblende) of sample F1. (e) Backscattered-electron image showing
two different phases of amphibole, where the high temperature hornblende was replaced by low temperature actinolite. (f) Garnet
porphyroblast of sample F93d. This porphyroblast has biotite, plagioclase, quartz and chlorite inclusions. (g) Deformed garnet crystal of
sample F93d with long axes parallel to the matrix foliation. (h) Cordierite porphyroblast of sample F93d with sigmoidal shape.
2009 Blackwell Publishing Ltd
clinopyroxene (in order of abundance) along with
minor titanite, magnetite and epidote. Clinopyroxene
occurs as anhedral elongated crystals up to
0.3 ·0.7 mm which are highly deformed and dyna-
mically recrystallized in the S
foliation (Fig. 4b).
Hornblende occurs as both small grains (<0.5 mm)
and porphyroblasts (>1.5 mm). The small crystals are
highly deformed, parallel to S
foliation and in equi-
librium with the clinopyroxene. The porphyroblasts
are subhedral to anhedral weakly deformed grains.
These porphyroblasts have inclusions of clinopyroxene
(Fig. 4c). The long axes of the hornblende porphyro-
blasts are sub-parallel to the S
foliation. The horn-
blende crystals are overgrown by actinolite. Plagioclase
occurs as subhedral to anhedral equant crystals
attaining 0.75 mm in diameter and highly altered to
epidote. Quartz with undulose extinction occurs
interstitially between the plagioclase, hornblende and
the clinopyroxene crystals. The magnetite occurs as
euhedral crystals (0.3 mm) that overgrow the S
ation. Titanite occurs as small euhedral inclusion
crystals (<0.01 lm) in hornblende porphyroblasts
within S
foliation (Fig. 4c).
Based on the observations above we interpret the
metamorphic peak assemblage as hornblende, clino-
pyroxene, plagioclase, magnetite and titanite and sug-
gest that it grew during D
or just thereafter. The
hornblende crystals are overgrown by actinolite and
epidote overgrows plagioclase. Therefore, the assem-
blage hornblende, actinolite, plagioclase, albite,
epidote and titanite is interpreted as a later assemblage
The mineral chemistry of the amphibole reveals
that they are magnesio-hornblende, edenite, edenitic-
hornblende, actinolitic-hornblende and actinolite
(Fig. 5a,b). The X
) ran-
ges (0.385–0.429), while the X
ranges around the
value 0.384. The Y(hb) = X(Al, M
) = (0.121–0.265)
and the Y(act) = (0.206–0.207) (Tables 1 & S1). The
cores of the plagioclase (pre-peak assemblage) have
X(an) = Ca (Ca + Na + K) ranges between 0.297
and 0.30, the plagioclase of the peak assemblage has
X(an) = (0.278–0.281), while plagioclase associated
with the late epidote and actinolite (post-peak assem-
blage), has X(an) = (0.252–0.278) (Table S2). Clino-
pyroxene has diopside–augite composition (Fig. 5c)
with X
) = (0.401–0.431)
(Tables 1 & S3).
Hornblende gneisses
The mineralogy of the hornblende gneisses includes
hornblende, actinolite, plagioclase, albite, K-feldspar,
quartz, chlorite as well as accessory amounts of
epidote, magnetite, ilmenite, titanite, apatite and
zircon. Plagioclase and amphibole are the most abun-
dant minerals, amounting to 80 vol.%. Plagioclase
occurs as anhedral to subhedral elongated crystals
attaining 0.35 ·0.7 mm. The plagioclase crystals show
albite and albite-Carlsbad twinning, but untwined
crystals are observed too. The plagioclase crystals are
highly retrogressed to epidote. Hornblende is a green
colour and strongly pleochroic. Hornblende occurs as
anhedral elongated crystals (1.63 ·0.68 mm)
(Fig. 4d). Hornblende is retrogressed to actinolite,
chlorite and titanite (Fig. 4e). Quartz occurs as anhe-
dral to subhedral crystals that lie interstitially between
the plagioclase and the hornblende crystals. It fre-
quently exhibits undulose extinction. The long axes of
the plagioclase and the hornblende crystals are parallel
to the metamorphic foliation (Fig. 4d). K-feldspar
occurs as very small (apparent only under the scanning
electron microscope) anhedral equant crystals, which
lie interstitially between the plagioclase, the quartz and
the hornblende crystals. Ilmenite grows over magnetite
as lamellar intergrowths.
The peak assemblage is identified by a continuous
metamorphic foliation, which is defined by parallel
crystals of hornblende, plagioclase, quartz and iron
oxides (Fig. 4d). The retrogression of plagioclase to
epidote and hornblende to actinolite, chlorite and
titanite as well as the presence of albite identify the
retrograde assemblage. The large difference between
the crystal size of the K-feldspar and the surrounding
minerals, the anhedral shape of the K-feldspar
crystals as well as the presence of K-feldspar-rich
metasomatic zones around the granitic dykes
(Fig. 3g) suggests that the K-feldspar are in disequi-
librium with the other phases.
Chemical analyses reveal that the studied amphiboles
are magnesio-hornblende, edenite, actinolitic-
hornblende and actinolite (Fig. 5a,b). The Al
ranges between 8.11 and 7.79 wt.% for the magnesio-
hornblende and the edenite, while this value ranges
between 2.69 and 4.34 wt.% for the actinolitic-
hornblende and the actinolite (Tables 1 & S4). The
ranges (0.475–0.477), while the X
(0.214–0.226). The Y(hb) = (0.063–0.066) and the
Y(act) = (0.16–0.22). The decrease in the Al
from earlier to later amphiboles (from hornblende
to actinolite) reveals that the amphibole goes from
relatively high temperature–high pressure field to low
temperature–low pressure field (Rasse, 1974; Blundy &
Holland, 1990) (Fig. 5d). The relationship between
and Si (Fig. 5e) indicates that the hornblende
crystallized under pressure of £5 kbar while the
pressure of the actinolite is most probably <5 kbar
(Fig. 5d). According to the graphical geothermometer
of Blundy & Holland (1990), the hornblende crystal-
lized at a temperature below 750 C, while actinolite
formed between 450–600 C (Fig. 5f).
In Wadi Aleiyat (Fig. 2) a few rock bands were found
that contain cordierite and garnet. These rocks have a
metagraywacke bulk composition similar to sample
ES356 of Sawyer (1986) which was modelled by
446 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
Johnson et al. (2008, e.g. their fig. 3a). Because these
rocks prove to be petrologically useful (see below),
they are described in detail here although they are
quite rare in the FSMC. The rocks contain biotite,
garnet, cordierite, plagioclase, quartz, chlorite, ilmen-
ite and rutile. Muscovite and K-feldspar are notably
absent. Garnet and cordierite occur as porphyroblasts
(Fig. 4f). Garnet porphyroblasts are subhedral to
anhedral equant grains up to 5 mm in diameter with
inclusions of quartz, plagioclase, biotite, chlorite and
ilmenite. The inclusions define linear trails that are
oblique to the matrix foliation (S
) giving relic evi-
dence to an earlier deformation phase prior to D
Some garnet is idiomorphic and has straight crystal
boundaries against biotite in the matrix. Other garnet
crystals are highly deformed with long axes parallel to
the matrix foliation (Fig. 4g). The cordierite por-
phyroblasts are anhedral, highly altered grains 0.5 mm
and 3 mm in diameter. The cordierite crystals are
equidimensional porphyroblasts with sigmoidal shape
Fig. 5. Amphibole chemistry. (a) and (b) Amphibole composition classified after Leake et al. (1997). (c) Ca–Mg–Fe composition
ranges and nomenclature of clinopyroxene (after Morimoto, 1988). (d) Shows the relation between the Al
and Al
in the
amphibole of samples F1 and F63. (e) Relation between Al
and Si of hornblende (after Rasse, 1974). (f) Al
v. temperature for
amphiboles after Blundy & Holland (1990). The open circle is the hornblende of the hornblende gneisses (the old phase), the filled
circle is actinolite of the hornblende gneisses (the young phase) and the rhomb is the amphibole of the migmatite. (g) Zoning
profile of XMg = Mg
) through garnet grain.
2009 Blackwell Publishing Ltd
in S
(Fig. 4h). Cordierite porphyroblasts have quartz,
plagioclase, biotite, chlorite and ilmenite inclusions.
Garnet and cordierite porphyroblasts are surrounded
by a matrix of biotite, plagioclase, chlorite, quartz,
ilmenite and rutile that define a continuous S
tion. The plagioclase occurs as euhedral to anhedral
elongate crystals attaining 0.15 ·0.3 mm. Quartz
occurs as anhedral to subhedral crystals that lie inter-
stitially between the plagioclase and the biotite crys-
tals. Biotite occurs as anhedral to subhedral flaky and
stumpy crystals attaining 0.25 ·0.75 mm and is
rarely overgrown by chlorite.
Garnet is weakly zoned (Fig. 5g). The cores of these
garnet crystals have Mg contents between 0.71 and
0.75 and Fe
contents between 1.91 and 1.92, with
) = (0.72–0.73) and Z(g) =
Ca (Fe + Mg + Ca) = (0.04–0.044). The rims of
these garnet crystals have Mg cation range between
0.61 and 0.69 and Fe
cation range between 1.83 and
1.9, with X
= (0.734–0.752) and Z(g) = (0.04–
0.048). The cordierite is Mg-rich, with Mg and Fe
cations (1.37–1.49) and (0.49–0.53) respectively. The
) ranges (0.250–0.261).
The biotite within the matrix (S
biotite), has X
) ranges (0.468–0.483) and
Y(bi) = X(Al, M
) = (0.27–0.29) (Table S5). The
chlorite minerals have X
) = (0.381–0.4) Table S6.
Biotite–plagioclase–garnet geothermometery of
Bhattacharya et al. (1992) gives a temperature of
668–707 C for the peak assemblage and 676–722 C
using the thermometer of Ganguly & Saxena (1984).
The same minerals gave a pressure range of 6.2–8 kbar
according to Hoisch (1990) (Tables 2 & S7). The cor-
dierite–garnet (garnet-rim) geothermometer according
to Thompson (1976) gave a temperature range between
654 and 649 C and a range between 635 and 638 C
according to Perchuk & LavrentÕeva (1983) Table S8.
The presence of highly deformed garnet crystals
which have long axes parallel to the matrix foliation S
indicates that it grew early during D
The sigmoidal
shapes of the cordierite crystals indicate that these
porphyroblasts grew later – possibly during D
. The
peak assemblage was identified by the presence of
garnet porphyroblasts. Based on these observations,
the peak metamorphism was interpreted as an event
between D
and D
. Biotite, chlorite, plagioclase,
quartz and ilmenite inclusions inside the garnet por-
phyroblasts were interpreted as the pre-peak assem-
blage while biotite, chlorite, plagioclase, quartz,
ilmenite, rutile (matrix mineral) with or without cor-
dierite were interpreted as the post-peak assemblage.
The presence of idiomorphic garnet crystals (Vernon,
2004) as well as the temperatures given by the biotite–
plagioclase–garnet geothermometer possibly indicates
that the rocks reached peak conditions above the
Table 1. Representative mineral analyses of samples F39, F1 and F63; amphibole (normalized to 23 O and ignoring H
O), plagioclase
(normalized to 8 O) and pyroxene (normalized to 6 O).
Sample no. F39 F1 F63
Assemblage hb di pl mt sph (peak assemblage) hb di pl sph (pre-peak ass.) hb act pl ep sph
(post-peak ass.)
hb act ab ep sph
(late post-peak
Mineral hb6 hb14 pl28 pl31 hb13 pl16 pl18 py2 py30 act2 pl3 pl15 ab3 ab8 hb3 act2 hb8 act3
47.28 48.78 61.35 61.61 44.89 60.9 61.04 50.06 48.4 49.44 61.96 61.65 68.59 69.03 42.56 54.09 42.72 52.35
0.58 0.47 0.03 b.d.l. 0.94 0.05 b.d.l. 0.12 0.96 0.57 b.d.l. 0.03 b.d.l. b.d.l 0.92 0.13 0.98 0.23
6.07 6.18 23.84 23.98 8.95 24.56 24.48 5.37 7.49 5.05 24.03 23.73 19.55 19.57 9.1 2.69 8.58 4.34
0.02 0.01 b.d.l. b.d.l. 0.05 0.03 b.d.l. 0.02 b.d.l. b.d.l. b.d.l. 0.03 0.02 b.d.l b.d.l 0.09 b.d.l. 0.06
FeO 16.85 16.1 0.27 0.15 18.45 0.29 0.26 14.54 16.81 15.27 0.24 0.42 b.d.l b.d.l 20.22 8.89 18.43 9.49
MnO 0.62 0.64 0.1 b.d.l 0.58 b.d.l 0.03 0.73 0.63 0.56 0.04 b.d.l. b.d.l b.d.l 0.69 0.56 0.45 0.49
MgO 12.66 13 0.04 b.d.l. 11.31 b.d.l. b.d.l. 12.18 11.66 13.5 b.d.l. b.d.l. 0.02 b.d.l 9.89 18.26 11.33 17.81
CaO 11.34 11.65 5.9 5.82 11.51 6.14 6.11 15.82 11.52 11.37 5.15 5.74 0.4 0.3 11.74 11.69 12.52 11.37
O 0.87 0.81 8.13 8.18 1.17 7.8 7.86 0.65 1.19 0.75 8.00 8.03 11.22 11.38 1.64 0.39 1.4 0.57
O 0.61 0.54 0.32 0.24 0.98 0.19 0.19 0.07 0.82 0.62 0.68 0.32 0.01 0.01 0.89 0.17 0.8 0.22
Total 96.94 98.22 99.97 99.98 98.87 99.96 99.96 99.56 99.48 96.98 100.12 99.95 99.8 100.29 97.65 96.96 97.21 96.93
O 2323 8 823 8 8 6 623 8 8 8 823232323
Si 7.05 7.15 2.73 2.73 6.64 2.7 2.71 1.74 1.83 7.32 2.74 2.74 2.99 3.00 6.44 7.71 6.52 7.49
Ti 0.06 0.05 0.00 – 0.1 0.00 – 0.31 0.02 0.06 0.00 – 0.14 0.01 0.15 0.02
Al 1.06 1.06 1.25 1.25 1.56 1.28 1.28 0.22 0.33 0.88 1.25 1.24 1.01 1.00 1.62 0.45 1.54 0.73
Cr 0.00 0.00 – 0.01 0.00 – 0.00 – 0.00 0.00 0.01 – 0.01
0.32 0.18 0.01 0.01 0.4 0.01 0.01 0.00 0.00 0.02 0.01 0.01 – 0.53 0.00 0.00 0.02
1.77 1.78 0.00 0.00 1.87 0.00 0.00 0.42 0.51 1.86 0.00 0.00 – 2.03 1.06 2.35 1.11
Mn 0.07 0.08 0.00 – 0.07 – 0.00 0.02 0.02 0.07 0.00 – 0.09 0.06 0.06 0.05
Mg 2.81 2.84 0.00 – 2.49 – 0.63 0.67 2.98 0.00 2.23 3.88 2.58 3.8
Ca 1.81 1.83 0.28 0.27 1.82 0.29 0.29 0.59 0.46 1.8 0.24 0.27 0.02 0.01 1.90 1.78 2.05 1.74
Na 0.25 0.23 0.7 0.7 0.33 0.67 0.67 0.04 0.08 0.21 0.68 0.69 0.95 0.96 0.48 0.1 0.41 0.15
K 0.11 0.1 0.02 0.01 0.18 0.01 0.01 0.00 0.04 0.11 0.04 0.02 0.00 0.00 0.17 0.03 0.16 0.04
0.387 0.385 0.429 0.384 0.475 0.214 0.477 0.226
Y(hb) or (act) 0.121 0.221 0.207 0.207 0.063 0.16 0.066 0.22
X(an) 0.281 0.278 0.300 0.297 0.252 0.278 0.020 0.010
0.401 0.431
Additional data are in the online data repository.
448 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
solidus. However, as the metagraywackes do not have
any partial melting indications in the field, we suggest
that the liquid mode in the rock during the peak con-
ditions was below 5% (Sawyer, 1999).
Pseudosections were constructed for several bulk
compositions to constrain the P,Tand Xo paths.
Four samples were chosen to represent the rock types
discussed above. Sample F39 was chosen to represent
the migmatites. It is from one of the pyroxene-bearing
mesosomes in the central part of the study area
(2842¢13¢¢N and 3339¢30¢¢E; Fig. 2). Samples F1 and
F63 are hornblende gneisses from the next higher
unit. They were collected at 2844¢28¢¢N and
3334¢44¢¢E (F1) and at 2843¢39¢¢N and 3332¢41¢¢E
(F63). Sample F93d is characterized by the presence
of garnet and cordierite porphyroblasts. This sample
was collected from one of the rare metagraywacke
bands which are intercalated with the hornblende
gneisses and the quartzofeldspathic gneisses at the
end of Wadi Aleiyat (2840¢52¢¢N and 3339¢37¢¢E;
Fig. 2).
For bulk rock chemical analysis a Bruker Pioneer S4
X-ray fluorescence spectrometer was used at the
Institute of Earth Science, Karl-Franzens-Universita
Graz, Austria. Samples were prepared as fused pellets
using Li
flux. The pseudosections were con-
structed using THERMOCALCTHERMOCALC tc330 (Powell & Holland,
1988) and the internally consistent dataset of Holland
& Powell (1998). The following a–x models were used:
amphibole (Diener et al., 2007); muscovite paragonite
(Coggon & Holland, 2002); biotite (White et al., 2007);
orthopyroxene (White et al., 2002); clinopyroxene
(Green et al., 2007); plagioclase K-feldspar (Holland
& Powell, 2003); garnet (White et al., 2007); cordierite
(Holland & Powell, 1998); chlorite (Mahar et al., 1997;
Holland et al., 1998); epidote (Holland & Powell,
1998); melt (White et al., 2007); ilmenite hematite
(White et al., 2000) and magnetite spinel (White et al.,
Migmatites P–T pseudosection
The XRF whole-rock analysis of sample F39 showed
that the K
O and the MnO constitute <1% of the
bulk rock (Table 3). Therefore, the P–T pseudosection
for this sample was calculated in the system
NCFMASHTO for the phases listed in the figure
caption of Fig. 6a and for quartz in excess. As there is
no melt model for this system, it is only meaningful to
interpret the sub-solidus parts of the P–T path for this
sample where water is likely to have been present.
Table 2. Representative mineral analyses and conventional thermobarometers of sample F93d; biotite (normalized to 11 O and
ignoring H
O), chlorite (normalized to 14 O and ignoring H
O), cordierite (normalized to 18 O and ignoring H
O), plagioclase
(normalized to 8 O) and garnet (normalized to 12 O).
Assemblage bi chl pl ru ilm H
O (latest post–assemblage) bi g pl ilm H
O (peak assemblage) bi cd pl ilm H
(post–peak assemblage)
Mineral bi11 bi20 chl4 chl5 chl31 chl33 g16 bi18 pl4 g17 bi17 pl5 cd3 cd8 bi3 bi13
35.33 25.20 25.76 25.34 28.13 27.93 37.37 35.46 59.93 38.05 36.02 63.68 48.91 47.47 34.63 34.13
1.78 2.14 0.38 0.38 0.12 0.26 0.17 2.49 b.d.l 0.07 1.61 0.09 b.d.l b.d.l 2.94 1.77
18.39 19.27 21.45 21.50 19.85 18.43 20.42 18.85 23.13 20.69 19.69 21.19 32.59 32.21 18.5 18.67
0.07 0.17 0.10 0.10 0.01 0.12 0.02 0.07 0.06 0.08 0.17 b.d.l b.d.l 0.07 0.26 0.04
FeO 19.20 18.64 20.35 20.17 21.96 21.92 30.07 19.36 0.94 30.08 19.20 0.59 5.79 5.92 18.12 19.19
MnO 0.16 0.16 0.11 0.11 0.23 0.23 3.14 0.13 0.03 3.32 0.16 0.05 0.29 0.11 0.22 0.48
MgO 12.21 11.18 17.99 17.83 18.60 18.54 6.29 10.32 1.66 6.08 10.68 0.46 9.7 9.38 11.41 12.21
CaO 0.12 b.d.l b.d.l b.d.l 0.01 0.09 1.30 0.12 4.42 1.43 0.13 7.7 0.06 0.04 0.13 0.16
O 0.41 0.42 b.d.l b.d.l b.d.l b.d.l b.d.l 0.41 9.17 b.d.l 0.42 7.20 b.d.l b.d.l 0.41 0.44
O 8.71 9.90 0.05 0.04 0.07 0.15 0.05 9.55 0.12 b.d.l 8.57 0.18 0.03 b.d.l 9.12 7.97
O 111114141414121181211818181111
Si 2.64 2.62 2.68 2.66 2.85 2.88 2.95 2.65 2.66 2.98 2.67 2.78 5.03 4.99 2.62 2.60
Ti 0.1 0.12 0.03 0.03 0.01 0.02 0.01 0.14 0.00 0.09 0.00 0.16 0.10
Al 1.62 1.69 2.63 2.66 2.37 2.24 1.9 1.66 1.21 1.91 1.72 1.09 3.95 3.99 1.65 1.67
Cr 0.00 0.01 0.01 0.01 0.00 0.01 0.00 0.00 0.00 0.01 0.01 0.01 0.01 0.002
0.00 0.00 0.00 0.00 0.00 0.00 0.13 0.00 0.07 0.1 0.00 0.04 0.00 0.00 0.00 0.00
1.2 1.16 1.77 1.77 1.86 1.89 1.9 1.21 0.00 1.9 1.19 0.00 0.49 0.52 1.14 1.22
Mn 0.01 0.01 0.01 0.01 0.02 0.02 0.21 0.01 0.00 0.22 0.01 0.00 0.03 0.01 0.01 0.03
Mg 1.36 1.24 2.79 2.79 2.81 2.85 0.75 1.15 0.11 0.73 1.18 0.03 1.49 1.47 1.28 1.38
Ca 0.01 – 0.00 0.01 0.11 0.01 0.21 0.12 0.01 0.36 0.01 0.01 0.01 0.01
Na 0.06 0.06 – ––––0.06 0.79 – 0.06 0.61 – 0.06 0.06
K 0.83 0.94 0.01 0.01 0.01 0.02 0.01 0.91 0.01 0.81 0.01 0.00 0.88 0.77
0.467 0.483 0.471 0.468
0.388 0.388 0.397 0.399
0.711 0.722
Z(g) 0.040 0.044
Y(bi) 0.26 0.31 0.31 0.39 0.27 0.29
Pkbar (Hoisch, 1990) 8.1 6.2
TC (Bhattacharya et al., 1992) 706 675
TC (Ganguly & Saxena, 1984) 723 697
2009 Blackwell Publishing Ltd
Thus, for Fig. 6a, water was chosen to be in excess. As
a reference for the onset of melting, we have super-
imposed the initial melting reactions in the system
NCKFMASHTO on Fig. 6a keeping in mind that all
equilibria at temperatures above this line must be
interpreted with care (as the water saturated assump-
tion does not hold there).
The P–T pseudosection for sample F39 is charac-
terized by a series of mineral assemblage fields with
steep boundaries (Fig. 6a). Only one small divariant
field (hb+act+di+pl+ab+ep+sph) appears at 7.4–
8.3 kbar and 520–535 C. The hb+di+pl+sph
pentivariant field is the largest field and is stable at
4.3–10 kbar and 525–795 C. Magnetite-in reactions
appear between 3 kbar – 485 C and 9 kbar – 800 C.
The proportion of magnetite increases with increasing
temperature. The amphibolite to granulite facies
boundary identified by the first appearance of ortho-
pyroxene, occurs at 760 C. The melting reactions
appear as near isothermal reactions between (625 C–
10 kbar) and (675 C – 3 kbar). The stability fields of
albite appear at low temperature conditions (430 C–
3 kbar) while this temperature increases with the
increase of pressure (560 C – 10 kbar).
The migmatite peak assemblage of hb+di+pl+
mt+sph appears on Fig. 6a in a quadrivariant field in
the region (580–760 C at 3 kbar), and narrows to
higher pressure where it terminates at 785 Cat
8.65 kbar. Within this field, the peak conditions can be
further constrained using the X
and the Y(hb)
isopleths. However, much of this field is within the
partial melting region for which Fig. 6a is not well
constrained. Thus, while isopleths indicate peak con-
ditions between 5.3 and 7.8 kbar we suggest that these
cannot be used for further interpretation. A lower limit
of the temperature at the metamorphic peak is identi-
fied as the first appearance of the melt at 660 C while
an upper temperature limit is identified by the absence
of orthopyroxene, which appears above 775 C.
Conditions for the pre-peak assemblage hb+di+
pl+sph are mainly defined on the basis of the different
chemistry of texturally earlier amphibole and plagio-
clase crystals. The estimated temperature for these
minerals is 555–655 C. Unfortunately, because of the
parallelism of the mineral isopleths (Fig. 6a), there are
no pressure constraints. The post-peak conditions are
well preserved by the presence of the hb+act+
pl+ep+sph assemblage. The hb+act+pl+ep+sph
assemblage was contoured for the X
and the X(an)
isopleths (Table 1), to constrain the retrograde pressure
and the temperature conditions. The isopleths reveal
that this mineral assemblage was stable at temperature
range of (455–520 C). Because of the presence of the
albite, the migmatites were re-equilibrated at low tem-
perature (<450 C). The pressure range at the low
temperature condition cannot be closely constrained,
but the actinolite chemistry (Fig. 5d) is consistent with
cooling most probably <5 kbar.
Hornblende gneiss P–T pseudosections
The P–T pseudosection for the hornblende gneisses
(samples F1 & F63) was calculated in the system
NCFMASHTO (Fig. 6b) for a bulk composition from
XRF neglecting 0.95 wt% K
O (Table 3). Calculations
were performed assuming water to be in excess which is
likely to be justified in the light of the fact that these
rocks show no evidence for partial melting.
The P–T pseudosection is characterized by a series
of mineral assemblage fields with steep boundaries,
except the garnet-in reactions (Fig. 6b). The garnet-
bearing assemblages appear at high temperature and
high pressure conditions (>600 C and >8.8 kbar).
Only one and small isothermal divariant field
(hb+chl+pl+ab+ep+sph+ilm) occurs at 3–8 kbar
and 457–520 C. All the assemblages below 565 C–
3 kbar and 620 C – 10 kbar are free of magnetite. The
first appearance of magnetite is in the trivariant
hb+chl+pl+g+ilm+mt and in the quadrivariant
hb+chl+pl+ilm+mt fields. The largest field in the
P–T pseudosection is the pentivariant field hb+pl+
ilm+mt. This assemblage is stable at the range of 572–
755 C at 3 kbar, 625 C 8.8 kbar and 791 C–
9.65 kbar. The albite-bearing assemblages appear at
low temperature conditions 460 C 3 kbar and
550 C – 10 kbar.
Table 3. Representative XRF analyses of different rock groups; major elements in wt%, trace elements (ppm).
Group Sample LOI SiO
MnO MgO CaO Na
Sum Ba Ce Cr Rb Sr) V Y Zn Zr
Hornblende gneisses F59 1.39 67.16 15.47 4.45 0.08 1.84 3.25 4.46 1.43 0.71 0.22 100.60 354 33 <20 39 483 86 22 86 160
F1 0.87 64.67 14.43 6.21 0.13 2.96 3.44 4.93 0.95 0.85 0.21 99.81 465 42 97 38 332 128 27 94 171
F63f 1.20 61.81 16.78 6.54 0.11 2.13 4.50 5.34 0.93 0.95 0.31 100.73 188 77 <20 <15 526 129 39 66 208
F4 0.54 62.47 15.34 6.68 0.15 3.07 5.00 4.96 0.94 0.97 0.22 100.47 253 54 105 <15 370 146 34 86 182
F7 0.27 66.93 15.11 6.57 0.07 1.80 2.40 4.71 1.70 0.89 0.17 100.81 679 57 61 39 377 116 32 59 227
Quartzofeldspathic gneisses F77 0.56 77.58 12.10 1.22 0.06 0.47 1.39 1.52 5.73 0.13 b.d.l 100.93 839 27 <20 195 31 <20 <20 48 108
F77a 0.65 81.62 10.21 1.12 0.13 0.43 2.20 2.30 1.77 0.11 b.d.l 100.67 720 33 <20 61 34 <20 24 24 103
F35 0.58 64.37 13.39 3.69 0.67 2.17 7.79 2.63 4.75 0.49 0.20 100.92 773 68 22 138 140 31 32 377 252
Calcsilicate F51b 13.18 40.59 4.20 0.93 0.22 1.36 37.23 1.26 0.64 0.17 0.07 99.90 115 <30 25 <15 371 50 <20 26 49
F33 1.33 53.83 9.91 4.09 0.28 2.07 23.87 2.63 1.92 0.40 0.16 100.58 349 43 30 42 133 66 26 205 108
Migmatite F39 0.75 59.91 14.59 6.20 0.24 4.08 8.13 4.21 0.73 0.78 0.21 99.96 167 60 96 <15 409 118 38 91 166
Metagraywackes F93d 1.57 60.82 14.60 8.91 0.11 5.00 2.01 2.34 3.15 0.97 0.15 99.83 483 84 429 90 162 143 25 96 202
F93a 1.00 46.95 2.17 34.93 0.07 2.21 7.44 b.d.l b.d.l 0.15 5.17 100.22 24 <30 26 <15 544 36 <20 61 88
F93c 1.00 47.72 4.26 27.53 0.09 2.31 10.05 0.25 b.d.l 0.22 6.83 100.41 83 <30 37 <15 574 43 <20 53 170
F94a 1.50 55.26 20.64 6.25 0.10 2.68 7.79 3.03 1.54 0.87 0.22 100.05 433 22 31 64 661 137 28 75 134
450 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
hb act pl sph
400 450 500 550 600 650 700 750
hb gl ab
ep sph
hb ab
ep sph
hb di ab
ep sph hb di pl
ab ep sph
hb di pl
ep sph
hb act ab
ep sph
hb act pl
ab ep sph
hb act pl
ep sph
hb act pl
mt sph
hb act di
ab ep sph
hb act di
pl ep sph
hb act di
pl mt sph
hb act di
pl sph
T (°C)
400 450 500 550 600 650 700 750
hb pl ilm mt
hb chl pl
g ilm
chl ab ilm
ep sp
ch l
ep sp h
chl pl ab ilm ep sp
chl pl
ab ep
sp h
Hb act
chl pl
hb act chl pl
ab ep sph
hb chl pl ab
ep sph ilm
hb chl ab
ep sph ilm
hb chl
ab ep
hb chl pl
ab ep ilm
hb chl pl
ep ilm
hb chl pl
ep sph ilm
NCFMASHT O system (+ q, H2O)
hb chl pl ilm mt
Hb pl
hb pl g ilm mt
hb pl
hb chl pl
g ilm mt
hb chl pl ilm
opx out
g out
act out
hb act ch
pl sph ilm
T (°C)
NCFMASHT O system (+ q, H2O)
Hornblende gneisses
ab out
0. 3
act chl pl ab ep sp
Hb chl pl ab ilm sp h
hb chl pl sph ilm
hb chl pl ab ep sp
Y (hb)
hb di pl
mt sph
hb di opx
pl sph
Y (hb)
X (an)
Y (act)
Transparen tr egio no f
pseudosection metastable
with respect to melt
hb di pl sph
Fig. 6. P–T pseudosections for migmatites
and hornblende gneisses. For both rock
types, the oxidation state was calculated
using Le Maitre (1976) equation for the
partitioning of FeO and Fe
. (a) P–T pseudosection of migma-
tite sample F39. The bulk composition is
(mol.%): SiO
: 64.12, Al
: 9.20, CaO:
9.38, MgO: 6.55, FeO: 5.41, Na
O: 4.38,
: 0.63, O: 0.32. The thick dashed line
shows the first appearance of the melt and
was calculatedin the system NCKFMASHTO
using K
O: 0.73%. As this line was not
calculated in the system for which the
pseudosection is drawn, the region above
it cannot be interpreted (shaded region).
(b) P–T pseudosection of average composi-
tion of samples F1 and F63. The bulk com-
position is (mol. %) SiO
: 69.56, Al
: 9.98,
CaO: 4.36, MgO: 4.22, FeO: 5.34, Na
5.53, TiO
: 0.75, O: 0.27. The vertical and
the horizontal bars are the average Pand T
according to Fig. 5. The white polygons
indicate conditions constrained by mineral
2009 Blackwell Publishing Ltd
The absence of orthopyroxene and garnet in the
peak assemblage hb+pl+ilm+mt constrains the peak
conditions to below 760 C and 9 kbar. X
Y(hb) contours (Table 1) were added to the peak
assemblage field in order to constrain the P–T condi-
tions of the equilibrium assemblage in more detail.
These conditions are agreement with the metamorphic
conditions derived from the relationship between
and the Si and the graphical geothermometer
(Fig. 5e,f). The post-peak conditions are difficult to
determine. Some constraints however are provided by
the late stage minerals chlorite, albite, actinolite
and epidote. The presence of these phases implies that
the hornblende gneisses re-equilibrated below the
albite-in boundary and above the actinolite-out
boundary (465–480 C and <<5 kbar according to
Fig. 4d,e).
Metagraywacke P–T pseudosection
In contrast to the migmatites and the hornblende
gneisses, the metagraywackes contain a series of Fe-
bearing phases (garnet, cordierite and biotite). For the
metagraywackes, we did not rely on Le Maitre (1976)
equation to constrain the oxidation state, but con-
strained it in more detail using a T–Xo pseudosection.
Indeed, the mineralogy of this sample does allow a
much closer constraint using this method than the
mineralogy of the other rock types would. A T–Xo
pseudosection (Fig. 7a) of the metagraywackes (sam-
ple F93d) was constructed in the system
NCKFMASHTO at 8 kbar (the pressure given by
biotite–plagioclase–garnet equilibria, Table 2). The
loss of ignition (LOI = 1.57) was used to evaluate the
O content of this sample. Z(g) isopleths and liquid
modes were added to the peak assemblage bi+g+
pl+ilm+liq in order to constrain the correct T–Xo
conditions. The Z(g) isopleths (0.04–0.044) intersected
with 0.05 liquid mode below 735 C and in a range of
Xo of 0.09–0.17. From the T–Xo pseudosection, the
best Xo estimate is 0.13, which is equivalent to an
oxidation state with a free oxygen content of 0.43.
Using this bulk composition, a P–T pseudosection
of the metagraywackes was constructed in the system
NCKFMASHTO. This pseudosection shows that
gedrite is stable over a wide P–T range – which is not
observed in the metagraywackes of the FSMC. We
suggest that this can be explained through a fraction-
ation in bulk composition during the retrograde path
¨we, 1997): garnet crystals are only zoned along
20% of their margins (Fig. 5g) indicating that most
of their volume was effectively removed from the bulk
composition at the peak. We have therefore used the
bulk composition described above only to interpret
the peak conditions (Fig. 7b). For the interpretation of
the retrograde evolution a new effective bulk compo-
sition was calculated by removing 80% of the garnet
from the bulk of Fig. 7b (Fig. 7c).
Both pseudosections (Fig. 7b,c) are characterized by
the absence of univariant and divariant assemblages.
There is one trivariant field; the bi+hl+pl+ep+
O at low temperature conditions.
The garnet-in reactions appear at high pressure con-
ditions (>6.5 kbar) in the temperature range 525–
720 C. Gedrite-bearing fields appear in the pressure
range 5–9.6 kbar and temperature range 600–790 C.
The cordierite-bearing assemblages appear at low
pressure conditions and high temperature (>600 C)
conditions. The orthopyroxene-in reactions appear as
isothermal boundaries at 765 C. At high pressure
conditions (>6.5 kbar), the orthopyroxene-in reac-
tions appears at >800 C. The melting reactions
appear as isothermal reactions (700–720 C). The
chlorite-bearing fields are stable at different pressure
conditions but at low temperature (<590 C). On the
other hand, albite is stable at the same pressure as
chlorite but at much lower temperature (<410–
530 C).
The metamorphic peak is constrained by the
assemblage bi+g+pl+ilm+liq on Fig. 7b. The
,Z(g) isopleths and liquid modes were added to
the field of the peak assemblage to determine the
peak conditions. The Z(g) is 0.04–0.0445 implying
that the peak pressure is 7.7–8.2 kbar. The X
ranges between 0.71 and 0.72, this indicates that
sample F93d reached peak conditions at 690–715 C.
This is consistent with P–T estimates obtained from
conventional thermobarometers as discussed above.
The intersections of the X
and Z(g) isopleths
support the assumption that the liquid mode of the
rock is not more than 5%. The presence of cordierite
and the absence of orthopyroxene and gedrite indi-
cate that sample F93d re-equilibrated at low pres-
sure–high temperature conditions (<6 kbar, 600–
760 C). Unfortunately, mineral isopleths of the
assemblage bi+cd+pl+ilm+H
O are more or less
isothermal which means that the temperature can be
determined. Y(bi) reveals that sample F93d re-equili-
brated at a temperature range 650–700 C. Low
temperature retrograde conditions could be identified by
the presence of late chlorite and rutile and the absence of
epidote and albite. The compositions of the late chlorite
and late biotite constrains the retrograde P–T condi-
tions to 3.5–4.8 kbar and 465–500 C (Fig. 7c).
The pseudosections discussed above constrain different
metamorphic conditions in different rock types (Figs 6
& 7). In this section, we discuss if these different
conditions reflect separate events or whether they can
be connected to a path for a single orogenic cycle. This
path is then related to the deformation history and an
integrated P–T–D evolution for the FSMC is derived.
It is also discussed as to why different rock types
record different stages of the evolution.
452 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
ged bi g
pl ilm liq
ged bi g
pl ilm
bi g pl
ged bi pl
ged bi chl
pl ilm H2O
ged bi chl g
pl ilm H2O
bi chl pl
bi chl ep ab
ru ilm
bi chl pl ep
ru ilm
bi chl g pl
ep ilm H2O
bi chl g pl
ged bi cd opx
pl ilm liq
bi cd pl
ilm liq
bi cd pl ilm
ged bi pl
ilm H O
bi cd opx pl ilm liq
ged bi cd
pl ilm liq
bi pl ru ilm
ged bi pl ilm liq
bi pl ilm
bi chl pl ru ilm
bi chl ep ab
ru ilm
bi chl pl ep
ru ilm H
400 450 500 550 600 650 700 750
NCKFMASHT O system (+ q)
T (°C)
NCKFMASHT O system (+ q) at
ar P8
X o
T (°C)
0.06 0.16 0.26 0.36
ged bi ilm
chl pl
ged bi ilm pl
ged bi g
ged bi g
liq ged bi ilm pl liq
bi g ilm pl liq
0.045 0.039
ged bi ilm liq pl H O
Liq mode
bi ilm chl pl
ged bi
pl ilm liq
bi g pl
ilm liq
0.04 0.05 0.06
0.7 0.71 0.72
Z (g)
liq mode
ged bi cd
pl ilm H O
Y (bi)
bi chl pl ep
bi chl pl ru
ilm H
Fig. 7. P–T and T–Xo pseudosections for
the metagraywackes (sample F93d).
(a) T–Xo pseudosection. The bulk com-
position is that derived from XRF analysis
(mol.%): SiO
: 66.41, Al
: 9.4, CaO: 2.35,
MgO: 8.14, FeO: 8.14, Na
O: 2.48, K
2.19, TiO
: 0.79, H
O: 5.41 (Table 3). (b)
P–T pseudosection at high pressure condi-
tions using the bulk composition derived
from (a) with Xo: 0.13. (c) P–T pseudosec-
tion at low pressure conditions using the
bulk composition of (b) from which 80% of
the garnet was removed (mol.%; SiO
: 64.35,
: 8.48, CaO: 2.25, MgO: 7.5, FeO:
5.04, Na
O: 2.53, K
O: 2.23, TiO
: 0.81,
O: 5.51). The vertical and the horizontal
bars are the average Pand Taccording to
the garnet–biotite–plagioclase thermo-
barometers (Table 2) and the garnet–
cordierite thermometer.
2009 Blackwell Publishing Ltd
P–T path and its relationship to deformation
In summary from above, pre-metamorphic peak
conditions are constrained by the chemistry of
clinopyroxene and plagioclase (and the absence of
magnetite) in the migmatites. These pre-peak condi-
tions are 550–620 C and 7.2–8.2 kbar. Peak meta-
morphic conditions derived for the metagraywackes
are 690–715 C and 7.7–8.2 kbar (Fig. 7). In the
migmatites, the peak is constrained at a similar pres-
sure, but temperatures 660–775 C (Fig. 6a). The
somewhat higher peak temperatures may simply reflect
that the migmatites (located in the core of the FSMC)
were located at slightly deeper crustal levels at the time
of peak metamorphism. However, it is also important
to note that the peak conditions of the migmatites were
derived without considering a melt model for this bulk
composition so that they are subject to some uncertainly.
Peak assemblages in both the migmatites and the
metagraywackes are overprinted by later parageneses
that equilibrated at 645–700 C and 4.7–5.3 kbar.
Finally, all studied samples give petrographical and
mineral chemical indications for partial re-equilibra-
tion at low pressure–low temperature conditions of
460–530 C and 4–5 kbar.
Interestingly, the metamorphic peak recorded by
the migmatites and metagraywackes is not preserved
in the hornblende gneisses. The coarse-grained peak
assemblage in the hornblende gneisses (hb+pl+
ilm+mt) formed at 630–770 C and <5.3 kbar.
These conditions correspond to the post-peak
assemblage in the other rock types. The final isobaric
cooling is recorded in all rock types. We suggest that
the migmatites and hornblende gneisses record dif-
ferent stages of the P–T path and interpret this as a
reflection of the water content of the rocks: in the
migmatites, water was partitioned into the melt at
the metamorphic peak, causing a relative dry residue
assemblage and thus preservation of the peak
assemblage (e.g. White & Powell, 2002). Conversely,
the paragneisses around the granitic dykes are en-
riched in K-feldspar which indicate metasomatic
infiltration. In addition all the metamorphic rocks
around the migmatites were affected by complex
fluid-infiltration processes after the peak metamor-
phism (Abu-Alam & Stu
¨we, Unpublished data). We
interpret therefore that the hornblende gneisses had
sufficient water during their entire evolution so that
later equilibration along the same P–T path of the
migmatites and metagraywackes was possible (e.g.
Guiraud et al., 2001).
The different metamorphic conditions described
here were interpreted by El-Shafei & Kusky (2003) as
evidence for two discrete metamorphic events: a
medium grade event M
(recorded by the hornblende
gneisses) and a high-grade M
event associated with
the formation of migmatites. However, Fowler &
Hassan (2008) suggested only one metamorphic event.
Textural, microstructural and thermobarometric evi-
dence presented here is agreement with Fowler &
Hassan (2008) and suggests that minor grade varia-
tions can be interpreted in terms of structural level and
different rock types. We infer therefore that the pre-
peak-, peak- and retrograde conditions reflect the
partial record of a single continuous ÔclockwiseÕP–T
path. This interpretation is consistent with geochro-
nological and tectonic evidence from elsewhere in the
Arabian-Nubian shield (Abu El-Enen et al., 2004;
Eliwa et al., 2008). Then, the inferred P–T path is
characterized by heating with moderate burial at mid
crustal levels (for an overburden density 2850 kg m
and assuming lithostatic conditions the peak meta-
morphic conditions corresponds to a depth of
25.7–29.3 km) followed by isothermal decompression
because of rapid exhumation to 5 kbar and final
cooling to stable geothermal conditions at this depth
(14.3–17.9 km). The final isobaric cooling path is
consistent with observations in Taba metamorphic
complex (the nearest high-grade metamorphic com-
plex) where final cooling is also near isobaric at 4.5–
5 kbar (Abu El-Enen et al., 2004). Exhumation from
this depth (14.3–17.9 km) to the surface is not
recorded by the metamorphic parageneses and will be
interpreted below.
The P–T evolution can now be correlated with the
four deformation phases that affected the Feiran–Solaf
region (Fig. 8). All peak parageneses grew in the S
foliation as well as the presence of leucosomes which
are parallel and cross-cutting to the principal foliation
suggesting that peak metamorphism was syn- or just
(Fig. 8). This interpretation is in agreement
with the observations of Fowler & Hassan (2008) in the
Feiran complex, but we note that it is different from
other metamorphic complexes in Sinai: in Taba and
Elat metamorphic complexes, Cosca et al. (1999)
suggested that the peak metamorphism was associated
with three ductile deformation phases D
, while
Abu El-Enen et al. (2004) suggested that the peak
metamorphism associated with D
However, we note that D
in Taba may correspond
to D
in the FSMC as we have found sparse evidence
for a deformation phase prior to peak metamorphism.
Microstructural evidence and field evidence shows that
and D
are compressive deformation phases. Thus,
these deformation phases must have occurred during
the isothermal decompression and or subsequent
cooling. As D
and D
structures are in places cut by
leucosomes, these phases are likely to have occurred
during the decompression phase above the solidus
(Fig. 8). In the Taba, Kid and Elat metamorphic
complexes, the D
and D
phases are also recorded as
ductile phases (e.g. Cosca et al., 1999; Abu El-Enen
et al., 2003, 2004; Eliwa et al., 2008) but their
relationship to the decompression path has not been
documented. On D
there are no direct constraints. We
suggest that it either occurred during cooling or during
the exhumation post-date to the 4 kbar cooling docu-
mented here.
454 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
The interpreted P–T–D evolution raises a series of
problems of both conceptual and regional relevance
that are now discussed. In particular, we discuss (i) the
counter-intuitive synchronous occurrence of isother-
mal decompression (recorded by the parageneses and
suggested exhumation) and convergent deformation
phases (recorded by field evidence and suggested bur-
ial) and (ii) the tectonic relevance of the documented
Exhumation during compressive deformation
The observation of synchronous decompression and
convergent deformation phases is not unique to the
FSMC. Carson et al. (1997) observed a similar rela-
tionship in the Larsemann Hills in Antarctica and in
several other high-grade complexes convergence is
observed synchronously with the decompression phase
of the P–T path. For the Larsemann Hills, Carson
et al. (1997) suggested that it reflects rapid upper
crustal extension at the time of lower crustal conver-
gence. However, for the FSMC the model of Carson
et al. (1997) is unsatisfactory, as the region is a rela-
tively small area for which orogen scale considerations
may not apply. In a more simple, one-dimensional
interpretation, Stu
¨we & Barr (1998) showed that
simultaneous erosion and convergence will result in a
deformation field characterized by decompression
during convergence above a critical point in the crust
and burial below that. This critical point separates an
upper part of the crust where upwards advection by
erosion outweighs convergence from a lower part
where deformation outweighs erosion. The depth of
this critical point depends on the relative rates of
erosion and deformation but can be as deep as
30 km. Because the FSMC records decompression
from 29 km, this model may, in principal, be a pos-
sible explanation. However, here we suggest that the
synchronous occurrence of convergent deformation
and decompression is simply a reflection of a compli-
cated obliquely convergent geometry of the D
and D
events which caused exhumation in a wrench system.
Within this wrench system overall convergence was
accompanied by extrusion of the complex to the sur-
face. Such a wrench system is well known from the
Pan-African basement in Sinai and the Eastern Desert:
the Najd fault system.
Tectonic implications in the context of the Najd fault system
The lithosphere-scale Najd fault system (Fig. 1) is
known to be related to the exhumation history of a
series of high-grade metamorphic terranes throughout
the Eastern Desert of Egypt and Sinai that are of
similar age and grade to the FSMC (615–610 Ma;
Stern & Manton, 1987): Pan-African metamorphic
areas are known for example in Wadi Hafafit (620–
580 Ma, Abd El-Naby et al., 2008; Loizenbauer et al.,
2001; Stern & Hedge, 1985; Hashad et al., 1981), Taba
complex (605 Ma, Eliwa et al., 2008), Meatiq meta-
morphic core complex (640–580 Ma, Fritz et al., 1996,
2002; Loizenbauer et al., 2001; Stern & Hedge, 1985)
and Sibai dome (614 Ma, Fritz et al., 1996, 2002). All
of these regions are characterized by low pressure–high
temperature peak conditions 6–8.5 kbar and 600–
750 C comparable with those described here (e.g.
Loizenbauer et al., 2001; Abd El-Naby et al., 2008;
Eliwa et al., 2008). Interestingly, most of the terranes
in the Eastern Desert were exhumed as core complexes
in extensional environments, while we interpret the
FSMC to have exhumed in an overall regime of
Here, it is argued that both mechanisms are possible
within the overall setting of the Najd fault system: this
fault system has a general NW–SE orientation (with
substantial local variation, see e.g. Shalaby et al., 2006)
and a sinistral motion, so that the maximum com-
pressive stress is oriented roughly W–E. We suggest
that minor variation in the orientation of the various
metamorphic complexes with respect to the Najd fault
system boundaries cause the different exhumation
mechanisms: in the Eastern Desert, the intermediate
principal stress may be oriented vertically causing
overall extension, while in the FSMC the minimum
principal stress is vertical causing exhumation in an
overall horizontal transpressive regime (Fig. 1). We
therefore suggest that the Najd fault system (D
in the
FSMC) is the principal phase causing the exhumation
of the complex in an oblique transpressive regime
(Fig. 9). Within this setting, major volumes of post-
tectonic granites that surround the complex were
produced by decompression-enhanced melting and are
transported through the Najd fault system, thereby
400 450 500 550 600 650 700 750
T (°C)
D1deformation phase and foliation.S1
23 4 3
, (compressional phases) and (open warping event),
formed open folds ( ) causing the map scale structure. At the end
of , the study area affected by the Najd Fault system.
Hornblende gneisses
Migmatite solidous
(Fig. 5)
610 – 615 Ma
Fig. 8. Summary of the P–T paths of metagraywackes,
hornblende gneisses and migmatites. The thick arrows mark
the P–T paths. The lines and the polygons are the stability fields
of the mineral assemblages from Figs 6 & 7.
2009 Blackwell Publishing Ltd
obliterating the margins of the complex. Our inter-
pretation is nicely consistent with interpretations from
the Eastern Desert where the D
phase is also inter-
preted as the Najd Fault system and related to ascent
of the post-tectonic granitic magma (e.g. Shalaby
et al., 2005; Farahat et al., 2007).
Final exhumation to the surface
An interesting and as yet unresolved aspect of the
tectonic evolution of the FSMC is its final exhumation
to the surface. This study has inferred a metamorphic
evolution that terminates at 450 C and 15 km
depth suggesting that the final exhumation must be
related to an independent event. Interestingly, this
depth is roughly the peak metamorphic depth of the
low-grade metamorphic belts in Sinai (Ayalon et al.,
1987; Eliwa et al., 2004) suggesting that these belts
were not affected by the Najd fault system. Zircon and
apatite fission track thermochronometers indicate that
the study area was subject to a complex history of
exhumation and reburial from the Cambrian to the
Early Tertiary period (e.g. Kohn et al., 1992). How-
ever, Vermeesch et al. (2009) suggested that the
calcalkaline rocks (at the same crustal level of the high-
grade gneisses) were exposed by c. 590 Ma suggesting
that the final exhumation occurred already at the end
of the Pan-African. In fact, in the FSMC, the presence
of an unconformity between gneisses and a Cretaceous
succession and the absence of the pre-Cretaceous
sedimentary cover (Fig. 2) suggest that the study area
was exhumed and subjected to intense pre-Cretaceous
erosion prior to being reburied under the Cretaceous
succession. This implies that the final exhumation of
the FSMC from 15 km occurred in an event fol-
lowing the Pan-African high-grade evolution. Such an
event has not been described, but there is sedimentary
evidence from the Hammamat molasse that subsidence
occurred and basins formed at c. 585 ± 15 Ma (Willis
et al., 1988). Within this event the lower-grade meta-
morphic belts of Sinai, like the SaÕal belt, may also
have exhumed.
In conclusion, the following tectonic evolution is
inferred from our study (Fig. 9). The rocks of the
Wadi Feiran belt include a series of metasedimentary
and metavolcanic rock types with minor mafic inter-
calations and calc-silicates that are consistent with an
interpretation as a sedimentary succession in a
marginal basin between Pan-African volcanic arcs
(El-Gaby & Ahmed, 1980). Vertical flattening (Fowler
& Hassan, 2008) was associated with horizontal
High grade
rocks (FSMB) Low grade
Volcanic arc
Ocean D
Najd system
(Extensional phase)
(Transpressive exhumation)
(Obliteration of contacts)
(Final exhumation)
Fig. 9. Tectono-metamorphic model of Wadi Feiran Solaf metamorphic complex during the Pan-African event. (a) Igneous
activity formed volcanic arcs and related intrusive rocks (632 ± 3) Ma. Immature sediments were deposited in a marginal basin,
followed by D
deformation due to vertical flattening. (b) Due to the arc–arc accretions and the closing of the Mozambique Ocean,
the D
and D
compressional phases occurred. (c) At the end of D
, the study area was exhumed by oblique compressive deformation
associated with sinistral strike–slip shear zone to the depth of 14.3–17.9 km. (d) The present day situation. The solid circles are the
position of the FSMC during the evolution.
456 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
transport towards the west and north-west, forming
the penetrative S
foliation and causing burial of the
rocks. The rocks reached peak metamorphism 7.2–
8.2 kbar and 645–775 C at 610–615 Ma towards the
end of D
and prior to D
. The subsequent D
and D
deformation phases caused shortening in NE–SW
direction at the time of substantial near isothermal
decompression of the rocks. The D
, phase is corre-
lated with the sinistral NW–SE striking Najd fault
system and exhumed the complex in an oblique
transpressive regime. We suggest that the difference
between the transpressive exhumation interpreted
here and the extensional core complexes of the
Eastern Desert is a consequence of minor variation of
the orientation of the various belts with respect to the
Najd fault system. Because of the rapid exhumation,
large volumes of post-tectonic granitic magma was
formed and transported along the Najd fault system
into the complex thereby obliterating most of its
margins. The doming up of the entire complex and
the NE–SW F
were formed by the open warping
event (D
) which was possibly related to the final
exhumation from 15 km depth to the surface during
an independent event.
This project was supported by the Austria exchange
service (O
¨AD) scholarship. We thank F. Makroum,
A. Shalaby, M. El-Shafei and Y.M. Sultan for their
help during the field work. Help with the pseudosec-
tions by V. Tenczer is appreciated. C. Hauzenberger is
thanked for his help with the XRF analysis. R.J. Stern
is thanked for his help with inaccessible papers.
S. Boger and M. Abu El-Enen are thanked for their
constructive reviews. R. White is appreciated for his
constructive suggestions and efficient editorial
handling of the manuscript.
Abd El-Naby, H., Frisch, W. & Siebel, W., 2008. Tectono-
metamorphic evolution of the Wadi Hafafit Culmination
(central Eastern Desert, Egypt). Implication for Neoprotero-
zoic core complex exhumation in NE Africa. Geologica Acta,
6, 293–312.
Abdelsalam, M. G. & Stern, R. J., 1996. Sutures and shear zones
in the Arabian-Nubian Shield. Journal of African Earth
Sciences,23, 289–310.
Abu El-Enen, M. M., Okrusch, M. & Will, T. M., 2003.
Metapelitic assemblages in the Umm Zariq schists, central
western Kid Belt, Sinai Peninsula, Egypt. Neues Jahrbuch
Mineralogie Abhandlungen,178, 277–306.
Abu El-Enen, M. M., Will, T. M. & Okrusch, M., 2004. P–T
evolution of the Pan-African Taba metamorphic belt, Sinai,
Egypt: constraints from metapelitic mineral assemblages.
Journal of African Earth Sciences,38, 59–78.
Ahmed, A. A., 1981. Reconsidered view on Feiran–Solaf
Gneisses southwest of Sinai, Egypt. Bulletin Faculty of
Sciences, Assiut University,10, 131–142.
Ahmed, A. A. & Youssef, M. M., 1976. Air photo interpretation
of dike swarms in the area around Feiran Oasis, southwest
Sinai, Egypt. Bulletin Faculty of Sciences, Assiut University,1,
Ayalon, A., Steinitz, G. & Starinsky, A., 1987. K–Ar and Rb–Sr
whole-rock ages reset during Pan-African event in the
Sinai Peninsula (Ataqa Area). Precambrian Research,37, 191–
Bhattacharya, A., Mohanty, L., Maji, A., Sen, S. K. & Raith,
M., 1992. Non-ideal mixing in the phlogopite–annite binary:
constraints from experimental data on Mg–Fe partitioning
and a reformulation of the biotite–garnet geothermometer.
Contributions to Mineralogy and Petrology,111, 87–93.
Blasband, B., Brooijmans, P., Dirks, P., Visser, W. & White, S.,
1997. A Pan-African core complex in the Sinai, Egypt. Geol-
ogie en Mijnbouw,76, 247–266.
Blasband, B., White, S., Brooijmans, P., De Brooder, H. &
Visser, W., 2000. Late Proterozoic extensional collapse in
Arabian-Nubian Shield. Journal of the Geological Society,
London,157, 615–628.
Blundy, J. & Holland, T., 1990. Calcic amphibole equilibria and
a new amphibole plagioclase geothermometer. Contributions
to Mineralogy and Petrology,104, 208–224.
Brooijmans, P., Blasband, B., White, S. H., Visser, W. J. &
Dirks, P., 2003. Geothermobarometric evidence for a meta-
morphic core complex in Sinai, Egypt. Precambrian Research,
123, 249–268.
Carson, C. J., Powell, R., Wilson, C. J. L. & Dirks, P. H. M. G.,
1997. Partial melting during tectonic exhumation of a granu-
lite terrane: an example from the Larsemann Hills, East
Antractica. Journal of Metamorphic Geology,15, 105–126.
Coggon, R. & Holland, T. J. B., 2002. Mixing properties of
phengitic micas and revised garnet–phengite thermobaro-
meters. Journal of Metamorphic Geology,20, 683–696.
Cosca, M. A., Shimron, A. & Caby, R., 1999. Late Precambrian
metamorphism and cooling in the Arabian-Nubian Shield:
petrology and
Ar geochronology of metamorphic
rocks of the Elat area (southern Israel). Precambrian Research,
98, 107–127.
Dachs, E., 2004. PET: petrological elementary tools for math-
ematica: an update. Computers and Geosciences,30, 173–182.
Diener, J. F. A., Powell, R., White, R. W. & Holland, T. J. B.,
2007. A new thermodynamic model for clino- and ortho-
amphiboles in the system Na
O–O. Journal of Metamorphic Geology,25, 631–656.
El-Gaby, S. & Ahmed, A. A., 1980. The Feiran–Solaf gneiss belt,
SW of Sinai, Egypt. In: Evolution and Mineralization of the
Arabian-Nubian Shield (eds Coory, P.G. & Tahoun, S.A.).
Institute of Applied Geology,4, 95–105, Jeddah.
El-Gaby, S., List, F. K. & Tehrani, R., 1990. The basement
complex of the Eastern Desert and Sinai. In: The Geology of
Egypt (ed. Rushdi, S.), pp. 175–184. Balkema, Rotterdam.
Eliwa, H. A., Abu El-Enen, M. M., Khalaf, I. & Itaya, T., 2004.
Metamorphic evolution of Sinai metapelites and gneisses:
constrains from petrology and K Ar age dating. Egyptian
Journal of Geology,48, 169–185.
Eliwa, H. A., Abu El-Enen, M. M., Khalaf, I. M., Itaya, T. &
Murata, M., 2008. Metamorphic evolution of Neoproterozoic
metapelites and gneisses in Sinai, Egypt: insights from
petrology, mineral chemistry and K–Ar age dating. Journal of
African Earth Sciences,51, 107–122.
El-Shafei, M. K. & Kusky, T. M., 2003. Structural and tectonic
evolution of the Neoproterozoic Feiran–Solaf metamorphic
belt, Sinai Peninsula: implications for the closure of the
Mozambique Ocean. Precambrian Research,123, 269–293.
El-Tokhi, M., 1992. Origin and tectonic implications of
Pan-African amphibolites of Wadi Feiran South Sinai. In:
Proceedings of the Third Conference on Geology of Sinai
Development (eds Abd El-Khalek, M. L., El-Ghawaby, M. A.
& Morsy, A. M.). Suez Canal University pp. 239–248. Ismailia.
Engel, A. E. J., Dixon, T. H. & Stern, R. J., 1980. Late
Precambrian evolution of Afro-Arabian crust from ocean to
craton. Geological Society America Bulletin,91, 699–706.
2009 Blackwell Publishing Ltd
Farahat, E. S., Mohamed, H. A., Ahmed, A. F. & El Mahallawi,
M. M., 2007. Origin of I- and A-type granitoids from the
Eastern Desert of Egypt: implications for crustal growth in the
northern Arabian-Nubian Shield. Journal of African Earth
Sciences,49, 43–58.
Fowler, A. & Hassan, I., 2008. Extensional tectonic origin
of gneissosity and related structures of the Feiran–Solaf
metamorphic belt, Sinai, Egypt. Precambrian Research,164,
Fritz, H., Wallbrecher, E., Khudeir, A. A., Abu El Ela, F. &
Dallmeyer, D. R., 1996. Formation of Neoproterozoic meta-
morphic core complexes during oblique convergence (Eastern
Desert, Egypt). Journal of African Earth Sciences,23,
Fritz, H., Dallmeyer, D. R., Wallbrecher, E. et al., 2002.
Neoproterozoic tectonothermal evolution of the Central
Eastern Desert, Egypt: a slow velocity tectonic process of core
complex exhumation. Journal of African Earth Sciences,34,
Ganguly, J. & Saxena, S. K., 1984. Mixing properties of
aluminosilicate garnets: constraints from natural and
experimental data, and applications to geothermo-barometry.
American Mineralogist,69, 88–97.
Gass, I. G., 1982. Upper Proterozoic (Pan-African) calcalkaline
magmatism in northeastern Africa and Arabia. In: Andesites
(ed. Thorpe, R.S.), pp. 591–609. Wiley, New York.
Green, E. C. R., Holland, T. J. B. & Powell, R., 2007. An order-
disorder model for omphacitic pyroxenes in the system
jadeite–diopside–hedenbergite–acmite, with applications to
eclogitic rocks. American Mineralogist,92, 1181–1189.
Guiraud, M., Powell, R. & Rebay, G., 2001. H
O in metamor-
phism and unexpected behaviour in the preservation of
metamorphic mineral assemblages. Journal of Metamorphic
Geology,19, 445–454.
Hashad, A. H., Sayyah, T. A. & El Manharawy, M. S., 1981.
Isotopic composition of strontium and origin of Wadi Kareim
volcanics, Eastern Desert. Egyptian Journal of Geology,25,
Hoisch, T., 1990. Empirical calibration of six geobarometers for
the mineral assemblage quartz+muscovite+biotite+plagio-
clase+garnet. Contributions to Mineralogy and Petrology,104,
Holland, T. J. B. & Powell, R., 1998. An internally consistent
thermodynamic dataset for phases of petrological interest.
Journal of Metamorphic Geology,16, 309–343.
Holland, T. J. B. & Powell, R., 2003. Activity–composition
relations for phases in petrological calculations: an asymmet-
ric multicomponent formulation. Contributions to Mineralogy
and Petrology,145, 492–501.
Holland, T. J. B., Baker, J. M. & Powell, R., 1998. Mixing
properties and activity–composition relationships of chlorites
in the system MgO–FeO–Al
O. European Journal
of Mineralogy,10, 395–406.
Johnson, P. R., 1998. Tectonic map of Saudi Arabia and
adjacents areas (scale: 1 4,000,000). In: Technical report
USGSOF-98-3 (ed. Saudi Arabian Deputy Ministry for
Mineral Resources), 2pp. Saudi Arabia, Jeddah.
Johnson, T. E., White, R. W. & Powell, R., 2008. Partial melting
of metagreywacke: a calculated mineral equilibria study.
Journal of Metamorphic Geology,26, 837–853.
Katz, O., Beyth, M., Miller, N. et al., 2004. A Late Neoprote-
rozoic (630 Ma) high-magnesium andesite suite from south
Israel: implications for the consolidation of Gondwanaland.
Earth and Planetary Science Letters,218, 475–490.
Khudeir, A. A., Abu El-Rus, M. A., El-Gaby, S., El-Nady, O. &
Bishara, W. W., 2008. Sr–Nd isotopes and geochemistry of the
infrastructural rocks in the Meatiq and Hafafit core
complexes, Eastern Desert, Egypt: evidence for involvement of
Pre-Neoproterozoic crust in the growth of Arabian-Nubian
Shield. Island Arc,17, 90–108.
Kohn, B. P., Eyal, M. & Feinstein, S., 1992. A major late
Devonian-early Carboniferous (Hercynian) thermotectonic
event at the new margin of the Arabian-Nubian Shield: evidence
from zircon fission track dating. Tectonics,11, 1018–1027.
¨ner, A., 1984. Late Precambrian plate tectonics and orogeny:
a need to redefine the term Pan-African. In: Ge
´ologie Africaine
(eds Klerkx, J. & Michot, J.), pp. 23–28. Muse
´e Royal de
l¢Afrique Centrale, Tervuren.
¨ner, A., Kru
¨ger, J. & Rashwan, A. A., 1994. Age and
tectonic setting of granitoid gneisses in the Eastern Desert of
Egypt and south-west Sinai. Geologische Rundschau,83,
Le Maitre, R. W., 1976. The chemical variability of some com-
mon igneous rocks. Journal of Petrology,17, 589–637.
Leake, B. E., Woollet, A. R., Arps, C. E. et al., 1997. Nomen-
clature of amphiboles: report of the subcommittee on
amphiboles of the International Mineralogical Association,
commission on new minerals and mineral names. American
Mineralogist,82, 1019–1037.
Loizenbauer, J., Wallbrecher, E., Fritz, H., Neumayr, P., Khu-
deir, A. A. & Kloetzli, U., 2001. Structural geology, single
zircon ages and fluid inclusion studies of the Meatiq
metamorphic core complex: implications for Neoproterozoic
tectonics in the Eastern Desert of Egypt. Precambrian
Research,110, 357–383.
Mahar, E. M., Baker, J. M., Powell, R., Holland, T. J. B. &
Howell, N., 1997. The effect of Mn on mineral stability in
metapelites. Journal of Metamorphic Geology,15, 223–238.
Morimoto, M., 1988. Nomenclature of pyroxenes. Mineralogical
Magazine,52, 535–550.
Perchuk, L. L. & LavrentÕeva, I. V., 1983. Experimental inves-
tigation of exchange equilibria in the system cordierite–
garnet–biotite. In: Kinetics and Equilibrium in Mineral
Reactions (ed. Saxena, S.K.), pp. 199–239. Springer Verlag,
New York.
Powell, R. & Holland, T. J. B., 1988. An internally consistent
thermodynamic dataset with uncertainties and correlations: 3.
Application, methods, work examples and a computer pro-
gram. Journal of Metamorphic Geology,6, 173–204.
Rasse, P., 1974. Al and Ti contents of hornblende, indicators of
pressure and temperature of regional metamorphism. Contri-
butions to Mineralogy and Petrology,45, 231–236.
Sawyer, E. W., 1986. The influence of source rock type, chemical
weathering and sorting on the geochemistry of clastic sedi-
ments from the Quetico metasedimentary belt, Superior
Province, Canada. Chemical Geology,55, 77–95.
Sawyer, E. W., 1999. Criteria for the recognition of partial
melting. Physics and Chemistry of the Earth,24, 269–279.
Shalaby, A., Stu
¨we, K., Makroum, F., Fritz, H., Kebede, T. &
¨tzli, U., 2005. The Wadi Mubarak belt, Eastern Desert: a
Neoproterozoic conjugate shear system in the Arabian-
Nubian Shield. Precambrian Research,136, 27–50.
Shalaby, A., Stu
¨we, K., Fritz, H. & Makroum, F., 2006. The El
Mayah molasse basin in the Eastern Desert of Egypt. Journal
of African Earth Sciences,45, 1–15.
Stern, R. J. & Hedge, C. E., 1985. Geochronologic and isotopic
constraints on Late Precambrian crustal evolution in the
Eastern Desert of Egypt. American Journal of Sciences,285,
Stern, R. J. & Manton, W. I., 1987. Age of Feiran basement
rocks, Sinai: implications for late Precambrian crustal evolu-
tion in the northern Arabian-Nubian Shield. Journal of the
Geological Society, London,144, 569–578.
Stern, R. J. & Manton, W. I., 1988. Discussion on the age of
Feiran basement rocks, Sinai: implications for late Pre-
cambrian crustal evolution in the northern Arabian-Nubian
Shield. Journal of the Geological Society, London,145, 1033–
¨we, K., 1997. Effective bulk composition changes due to
cooling: a model predicting complexities in retrograde reaction
textures. Contributions to Mineralogy and Petrology,129,
¨we, K. & Barr, T. D., 1998. On uplift and exhumation during
convergence. Tectonics,17, 80–88.
458 T. S. ABU-ALAM & K. STU
2009 Blackwell Publishing Ltd
Thompson, J. B., 1976. Mineral reactions in pelitic rocks; II,
Calculation of some P–T–X (Fe–Mg) phase relations.
American Journal of Science,276, 425–454.
Vermeesch, P., Avigad, D. & McWilliams, M. O., 2009. 500 Myr
of thermal history elucidated by multi-method detrital ther-
mochronology of North Gondwana Cambrian sandstone
(Eilat area, Israel). Geological Society of America Bulletin,121,
Vernon, R. H., 2004. A Practical Guide to Rock Microstructure.
Cambridge University Press, Cambridge, 594 pp.
White, R. W. & Powell, R., 2002. Melt loss and the preservation
of granulite facies mineral assemblages. Journal of Metamor-
phic Geology,20, 621–632.
White, R. W., Powell, R., Holland, T. J. B. & Worley, B. A.,
2000. The effect of TiO
and Fe
on metapelitic assemblages
at greenschist and amphibolite facies conditions: mineral
equilibria calculation in the system K
.Journal of Metamorphic Geology,18,
White, R. W., Powell, R. & Clarke, G. L., 2002. The interpre-
tation of reaction textures in Fe-rich metapelitic granulites of
the Musgrave Block, central Australia: constraints from
mineral equilibria calculations in the system K
.Journal of Metamorphic
Geology,20, 41–55.
White, R. W., Powell, R. & Holland, T. J. B., 2007.
Progress relating to calculation of partial melting equilibria
for metapelites. Journal of Metamorphic Geology,25, 511–527.
Willis, K. M., Stern, R. J. & Clauer, N., 1988. Age and geo-
chemistry of Late Precambrian sediments of the Hammamat
series from Northeastern Desert of Egypt. Precambrian
Research,42, 173–187.
Additional Supporting Information may be found in
the online version of this article:
Table S1. Representative amphibole analyses of
sample F39 (migmatite) (numbers of ions on the basis
of 23 O, ignoring H
Table S2. Representative plagioclase analyses of
sample F39 (migmatite). The chemical formula based
on 8 O.
Table S3. Representative cpx analyses of the assem-
blage (hb+di+pl+sph) sample F39 (migmatite). The
chemical formula based on 6 O.
Table S4. Representative amphibole analyses of
samples F1 and F63 (hornblende gneisses) (numbers of
ions on the basis of 23 O, ignoring H
Table S5. Representative biotite analyses of sample
F93 (metagraywacke) (numbers of ions on the basis of
11 O, ignoring H
Table S6. Representative chlorite analyses of sample
F93 (metagraywacke) (numbers of ions on the basis of
14 O, ignoring H
Table S7. Representative garnet, biotite and plagio-
clase analyses of sample F93 (metagraywacke). The
chemical formula based on, 11 O, ignoring H
for the biotite, 12 O for the garnet and 8 O for the
Table S8. Representative garnet and cordierite ana-
lyses of sample F93 (metagraywacke). The chemical
formula based on, 18 O, ignoring H
O for the
cordierite and 12 O for the garnet.
Please note: Wiley-Blackwell are not responsible for
the content or functionality of any supporting
materials supplied by the authors. Any queries (other
than missing material) should be directed to the
corresponding author for the article.
Received 31 October 2008; revision accepted 15 May 2009.
2009 Blackwell Publishing Ltd
... The postcollisional magmatism in southern Sinai was formed as a result of tectonic escape after completion of collision between continental plates and/ or island arcs, with associated regional metamorphism during orogenesis (e.g., Stern 1994;Kusky and Matsah 2003;Abu-Alam and Stüwe 2009). This stage was represented by vast intrusions of granites and related volcanic rocks (Farahat et al. 2007El-Bialy 2010;Eyal et al. 2010;Be'eri-Shlevin et al. 2011;Johnson et al. 2011). ...
Full-text available
Postcollisional magmatism is widely distributed in southern Sinai, the extreme northern part of the Neoproterozoic Arabian-Nubian Shield. This article deals with mineral and whole-rock chemistry of postcollisional syenogranites and associated volcanic rocks from three localities in southern Sinai: Iqna Sharay'a, Rusis-Rutig, and Um Shuki-Abu Khusheib. The studied volcanic rocks have compositions between rhyolites and dacites with minor andesite. The whole-rock chemical compositions of the investigated rock types together with the biotite chemistry are consistent with high-K calc-alkaline and alkaline/peralkaline magma. The studied syenogranites and most volcanic rocks are more akin to anorogenic alkaline within-plate environments. Only a few samples of Um Shuki-Abu Khusheib volcanic rocks display some characteristics of orogenic arc-type environments. The high-K calc-alkaline to alkaline affinity and the relative enrichments in large ion lithophile elements (especially K, Rb, and Ba) and light rare earth elements together with a significant negative Eu anomaly imply that the studied granites and volcanic rocks were generated by partial melting of lower to middle crustal materials accompanied by the underplated mafic magma produced in the lithospheric mantle (convective diffusion). This convective diffusion describes a specific scenario of active chemical interaction between mafic and silicic magmas in order to explain formation of voluminous high-K calc-alkaline and alkaline/peralkaline magmatism in postcollisional tectonic environments. The enhanced temperatures of A-type silicic magmas of more than 10007C suggest that magma generation could occur even at the depth of the uppermost lithospheric mantle.
... Plagioclase in the boninitic diabase is classified as albite, due to the effects of low-grade albitisation, whereas in the gabbro and the tholeiitic diabase it appears as albite, labradorite and andesine. Magmatic amphibole is present within the gabbro and to a smaller extent in the tholeiitic diabase as interstitial crystals and blebs, displaying lower silica and higher Ti, Al and Na contents compared to secondary amphiboles, such as tremolite and actinolite (Maeda et al., 2002;Koutsovitis and Magganas, 2016;Abu-Alam and Stüwe, 2009;Ridolfi et al., 2010). Based on the textural characteristics and the chemistry of amphiboles in the gabbros and tholeiitic diabasic rocks of our study, we consider that blebs and granular hornblende were crystallised during the main magmatic stage from a hydrous silicate melt, whereas fibrous amphibole (actinolite) formed at the expense of magmatic hornblende and clinopyroxene during the metasomatic stage. ...
In the Veria-Naousa ophiolitic complex (north Greece), rodingite appears mainly in the form of cross cutting dykes within serpentinised peridotites. It is distinguished into three types, based upon the provenance of its protoliths, textural characteristics, mineralogical assemblages and geochemical affinities. Type I rodigites were derived from boninitic diabasic protoliths and their mineralogical assemblage include garnet + clinopyroxene + chlorite. Type II rodingites were formed at the expense of gabbroic precursors, comprising clinopyroxene + garnet + vesuvianite ± quartz, whereas Type III rodingites replaced diabasic tholeiitic protoliths comprising of garnets + vesuvianite + clinopyroxene + chlorite. Rodingitisation resulted in desilification, decrease of alkalies, Al, Fe, Mg and increase in Ca contents. In Type I rodingites the MREE (middle rare earth elements) and HREE (heavy rare earth elements) were slightly reduced. Type II rodingites experienced LREE (light rare earth elements) depletions, whereas MREE and HREE remained fairly stable. Restricted mobility of REE in Type III rodingites is assigned to shallow-level rodingitisation under decreasing pH. Rodingitisation occured in two distinct stages at fore-arc settings. The first stage occured under mildly oxidising conditions and enhanced CO2/H2O ratios. This stage affected the protoliths of all rodingite types. The second rodingitisation stage occured under more oxidising conditions and lower CO2/H2O ratios, which corresponds to the exhumation stage of the serpentinite-rodingite formations. Types II and III rodingites were subjected to further rodingitisation under the increasing influence of slab-derived hydrous phases at shallower depths, leading to the formation of late-stage andradite and vesuvianite. All stages of rodingitisation are estimated to have occurred under relatively moderate temperatures and pressure (~300 to 450 °C; ~2–6 kbar respectively).
The “Supercontinents” constitute the majority of the ancient Earth’s surface and play an important role in Earth’s history.
This study concerns the structural setting of the central Egyptian-Nubian Shield (El Shalul area) utilizing field-structural, remote sensing, petrological and geochemical data. The exposed basement comprises ophiolitic-mélange, arc-related metavolcanics, metasediments, metagabbro-diorites and granitoids. The area experienced two stages of deformation, pre-Najd (∼850–630 Ma) and Najd-related (∼630–580 Ma). The pre-Najd stage is represented by the assembly of arc-terranes and their N-ward extrusion while the Najd-related stage encompasses three deformation phases. D1 is post-collision extensional event, depicting lateral spreading of tectonic terranes and NW-ejection of ophiolites. The emplacement of El Shalul granite (∼630 Ma) and deposition of molasse sediments in E-W and NW-SE extensional basins (Zeidun and Meesar) are D1- related. Following extension is a protracted phase of compression, shearing and transpression. It was commenced with NW-SE shortening (D2) and deformed the extensional basins by folding and thrusting with mild effect on the other basement units. Sinistral shearing (D3a) and transpression (D3b) along the NW-trending faults superseded the D2 compression, while dextral shearing (D3c) on the ENE-WSW and NE-SW faults overprinted the NW-SE penetrative structures and controlled the emplacement of post-granitic dykes. The significant conclusions of this study include (1) El Shalul granite complex is a large alkaline granitic sheet emplaced during a post-collision extensional regime and suffered the subsequent ∼630–580 Ma top-to-NW sinistral shearing and SW-directed thrusting (not a gneissic core complex) and (2) The NW-SE and NE-SW structural trends are not conjugated.
The present study deals with the geological and geophysical studies of El-Barramiya area between longitudes 33° 45 ′ E and 33° 55 ′ E and latitudes 25°00′ N and 25° 10′ N with focusing on the gold mine area (33° 47′ E and 25° 04′ N). Geological and structural setting, petrography, fluid inclusions and geochemistry of alteration and mineralized veins were carried out to determine the genesis of gold mineralization. Geophysical studies are including magnetic and gravity studies to delineate the surface and subsurface structures and shear zones which control gold mineralization.
Full-text available
The Neoproterozoic metavolcanic rocks constitute major rock sequences in the Egyptian Nubian Shield. A recent detailed overview of the volcanisms and volcanic sequences particularly the metavolcanic rock unit is treated in this chapter. The geologic overview is stressed on a brief description of the metavolcanics at 41 better studied occurrences, a brief description of some common petrographic characteristics of these rocks, a discussion of their metamorphism, geochemistry, and age-dating. Two essential types: ophiolitic metavolcanics (Oph Mv) and arc metavolcanics (Arc Mv) have been discussed in detail. They were overprinted by greenschist and lower amphi-bolite facies metamorphism. The ophiolitic (Old) metavol-canic rocks (Oph Mv) are concentrated in the Central Eastern Desert (CED) and Southern Eastern Desert (SED) of Egypt, comprise pillow lavas and sheeted dykes. The pillowed lavas display pillow structure with abundant amygdales and scarcely massive. The sheeted dykes' are repeated and continuous vertical styles. The rock types of the pillow lavas include spilite, aphyric and variolitic metabasalt, metabasaltic andesite and/or meta-andesite. The sheeted dykes exhibit subophitic, ophitic, and blastodiabasic metabasalt and rare metado-lerite and meta-andesite. The arc (Young) metavolcanic rocks (Arc Mv) are generally widespread in the southern Sinai, NED, CED, and northern parts of the SED of Egypt. They belong to the island-arc mafic to felsic metavolcanic rocks. The Arc Mv includes massive and schistose flow and/or pillow lavas, commonly associated with pyroclastic and tuffs. They comprise two associations of mafic and felsic: blastodiabasic and porphyritic metabasalt and meta-andesite association, and porphyritic meta-andesite, metarhyodacite and metadacite association. Geochemically, the Arc Mv has higher ranges and means of SiO 2 , Al 2 O 3 , CaO, Na 2 O, Sr, Nb, Zr, and V and lower MgO, MnO, TiO 2 , Ba, Rb, Cr, and Ni compared with that of Oph Mv, suggest their fractionation sources. On discrimination diagrams, transitional tectonic environment among MORB, within-plate, and island-arc basalt types of the Oph Mv, revealing simulating their back-arc marginal basin. However, N-MORB affinity for the lavas and fore-arc character for the associated ophiolitic metaultra-mafics have been also reported. These variations reflect source region heterogeneities, and differences in degree of partial melting as well as corroborating the variability of oceanic domain of the ED Mv. The complied REEs data can be grouped into: Oph Mv dominating in the CED and SED and Arc Mv of Southern Sinai, NED and CED. The Oph Mv consequently comprises two subgroups: (i) LEEs-depleted pattern (similar to the N-MORB-type) and (ii) LREEs-enriched pattern (analogous to transitional MORB-IA). On the other hand, the Arc Mv has REEs patterns similar to that of IA (e.g., LREEs-enriched). These variations also reflect source region heterogeneities, differences in degree of partial melting, and variability of oceanic-continental terrains domain of the Egyptian shield Mv. Bimodal arc associations of both mafic and felsic compositions (Shadli-type) are concentrated in CED and northern parts of the SED. The mafic end-members are tholeiitic basalt and basaltic andesite whereas the felsic ones comprise calc-alkaline dacite and rhyolite. The mafics are comparable to N-MORB and arc-related non-cumulative mafic rocks. Felsic end-members are subduction-related source of arc-related magmas of thickened low-K mafic lower crust. The Oph Mv is relatively yielded age ranges from 750-700 Ma, whereas, the Arc Mv is relatively range in ages from 750-610 Ma, suggesting an overlap and no discernible difference in age are found between both sequences. These ages may represent the time of eruption of these metavolcanics in the crustal growth of the Egyptian Nubian Shield.
The Saharan Metacraton is a poorly known tract of pre-Neoproterozoic continental crust that occupies the area between the juvenile Arabian Nubian Shield, in the east, and the Tuareg Shield to the west. Neoproterozoic orogenesis (i.e. Oubanguides and East African orogenies) affect the west, the south and the east of the metacraton, respectively, which led to deformation, emplacement of igneous bodies, and localised episodes of rift-related magamatism. Details about interior regions of the Saharan Metacraton are poorly known, with much of it covered by Phanerozoic rocks. The basement outcrops in Sudan, Chad, Algeria and Libya have been the subject of few modern geochronological studies. Here we present results from the first zircon geochronology and in-situ zircon hafnium isotope investigations from both the Sudanese Butana and central Chad. The terranes in Butana, formed to the east of pre-Neoproterozoic continental crust of the Saharan Metacraton, with the oldest juvenile magmatism (ƐHf(t) of +4.89 to +7.89) at ca. 839 Ma, followed by subsequent magmatism at ca. 787 Ma. The ca. 787 Ma event (seen elsewhere in the East African Orogen) is interpreted to represent volcanic-arc collision and accretion with the kernel of the Saharan Metacraton. The ca. 839 Ma magmatism is contemporaneous with the accretion of the Tonian (ca. 850 Ma) arc terranes of the Arabian Nubian Shield and marks subduction to the east of the Saharan Metacraton. The magmatic history of the Ouaddaï region in Chad begins in the late Mesoproterozoic, with localised rifting, resulting in the emplacement of juvenile granites (ca. 1030 Ma). The Cryogenian and Ediacaran in the Saharan Metacraton reveal a complicated history of magmatism and deformation. Age data from Chad show the emplacement of granites (from melting of Mesoproterozoic crust: ƐHf(t) = +2.04 and −4.07) at ca. 665–654 Ma, coeval with the main East African Orogeny and accretion of the other ANS terranes to the Saharan Metacraton at ca. 650–580 Ma. The youngest tectonothermal event within the Saharan Metacraton is recorded by emplacement of granites between 580 and 550 Ma (ƐHf(t) values of −17 to −31) in southern Chad.
The Arabian-Nubian Shield (ANS) was assembled from juvenile crust during a three-stage Neoproterozoic tectonic evolution involving: (1) intra-oceanic subduction and arc accretion stage, (2) orogenic extension stage, and (3) post-extensional compressional stage. Stage 1 is manifested by ophiolite-decorated arc-arc high-strain zones (suture zones) and calc-alkaline magmatism. The orogenic extensional tectonic stage generated dyke swarms, bimodal volcanism, molasse basins, A-type granite magmatism, low angle normal faulting (LANFs) and metamorphic gneiss complexes. The geological features attributed to this stage have been interpreted in terms of continental rifting, gravitational collapse, crustal and mantle delamination, transpression, escape/extrusion tectonism, and gravitational uplift. The post-extensional compressional stage is typified by dominantly NW–SE trending folds and thrusts, E–W transpression, and the N–S shortening zones. The Najd Fault System (NFS) (ca. 630–540 Ma) to be described in this chapter is attributed by some workers to the orogenic extension tectonic stage and by others to the post-extensional compressional stage. Earlier interpretation connects the NFS to the Najd Orogeny (570–520 Ma). The NFS is one of the largest transcurrent shear systems worldwide and deciphering its kinematic history adds considerably to our understanding of the cratonization of Gondwana, and specifically to mechanisms of exhumation of metamorphic complexes in the ANS. The NFS extends in a NW–SE direction across the Arabian Shield (e.g., Ajjaj, Qazaz, Ruwah, Ar Rikah, and Halaban strands of the NFS) for more than 1300 km (~400 km wide) and continues beneath Phanerozoic cover in Yemen. The NFS is believed to extend into the Nubian Shield (Egyptian Eastern Desert and Sinai). The dominant sense of shearing along the NW–SE trending Najd megashears is sinistral, however, evidence exists for an earlier phase of dextral slip. NE- (to ENE-) oriented shear zones (e.g. the Ad-Damm, Fatima, Idfu-Mersa Alam, Qena-Safaga shear zones) could be Najd-related conjugates or earlier fault systems. The shear and volume strain aspects of Najd shears are described, as are the stress controls on the brittle evolution of Najd faults. The role of Najd brittle structures in hydrothermal mineral deposits and ground water flow patterns are also covered in this chapter.
The Neoproterozoic East African Orogen (EAO) preserves one of the finest records of a complete Pre-Cambrian Wilson cycle that started with the fragmentation of Rodinia supercontinent (~870–800 Ma) and opening of the Mozambique Ocean, followed by convergence along easterly and/or westerly dipping subduction zones, and was completed by closure of the ocean basin and collision between a collage of continental fragments that comprised East and West Gondwana. A prolonged (~200 Ma) convergence culminated in arcs suturing and terranes accretion followed by arc-continent terminal collision to form a N–S oriented collision zone (EAO) differentiated into the Arabian-Nubian Shield (ANS) to the north and the Mozambique Belt to the south. Terrane accretion resulted in a substantial crustal thickening and differential uplift, of certain parts of the ANS, followed by erosion and deposition of thick cover sequences in a number of, mostly fault-controlled, depositional basins. The latter records a post-accretion complex array of tectonic—“active” and non-tectonic—“passive” related deformations. Active tectonics are attributed to the crustal shortening that accompanied terminal collision and wrenching whilst passive tectonics could be linked to extensional collapse of parts of the orogen and account for the low angle normal and reverse shear zones. Reactivation of pre-existing accretion-related lineaments, and the formation of wrench-related new shears might have created local and overlapping stress fields that resulted in variably oriented structures deviating drastically from the general stress field(s). The differential uplift across the ANS and the use of common criteria to interpret different tectonic regimes (e.g. extension versus wrenching), wherein the deformation events have overlapping, if not matching, dates, make the idea of a regional tectonic model of the ANS inapplicable. Moreover, in many parts of the shield, it is very plausible to interpret some post-accretion deformation events from different tectonic regimes perspective.
Compiled U–Pb and Pb–Pb zircon ages from the Sinai Peninsula Precambrian basement record a crustal history covering of a time span of ~450 Ma, which reflects a complex history of magmatism, sedimentation, and metamorphism extending from ca. 1030–578 Ma. Three magmatic main age populations are recognized at 1025–975, 850–725, and 650–575 Ma and are interpreted to represent three discrete calc-alkaline volcanic arc magmatic episodes for the Sinai basement, succeeded by a phase of post-collisional within-plate alkaline-peralkaline magmatism at the end of the Neoproterozoic time. Detrital zircon ages show three age populations comparable to those of the magmatic zircons, suggesting that the host sedimentary rocks were deposited within basins adjacent to the volcanic arcs. Provenances of the sediments include Neoproterozoic arc assemblages, and limited contributions from old continental crust trapped within the Arabian-Nubian Shield or located at its border with the Saharan Metacraton. Metamorphic zircons within the Sinai metamorphic rocks give ages consistent with two discrete high-grade metamorphic events at 800–750 Ma and 627–592 Ma in the Kid and Feiran-Solaf metamorphic complexes, respectively. The latter metamorphic event is divisible into two sub-events: a regional mid-crustal level upper amphibolite facies event, and a shallow-crustal level isothermal decompression event, caused by heat released from the youngest arc magmatism in the Sinai. Inherited zircons of Paleoproterozoic, Mesoproterozoic, and older ages are infrequent in the magmatic rocks of the Sinai basement. The common age of zircon xenocrysts is Neoproterozoic, implying a low contribution of reworked older crust.
Full-text available
Folded migmatitic gneisses and plutonic rocks occupy the area around Feiran Oasis. They are dissected by dyke swarms of different types and generations. The most abundant are older micro-granitic dykes (77.6^ of the total number). Ne¬arly all dykes trend in the NE - SW direction except two long olivine dolerite dykes which have WBW - ESE trend. The dykes are affected by faults except the olivine dolerite ones . The trend of most of the dykes, seems to be related to deep tensional fractures developed during a folding process.
The Feiran-Solaf gneiss belt represents a thick sedimentary succession with minor basic magmatic intercalations that was folded into three anticlines separated by thrust faults and metamorphosed into gneisses and migmatites. The fold axes and thrust faults trend nearly N.W.-S.E., The gneisses and migmatites are grouped into five formations. The western Feiran anticline comprises two formations: a lower semipelitic Nidia El-Samra Formation, and an upper calcareous-pelitic Aleiyat Formation. The central Solaf anticline also comprises two formations: and a lower pelitic and calcareous-pelitic El-Khali’ Formation, and an upper arenaceous Um Tarr Formation. The calcareous-pelitic El-Sheikh Formation forms the western limb of an incomplete and overturned Sheik anticline thrust over the Solaf anticline from the east. The metasomatic changes which the paragneisses have suffered are also considered.