The Mantle Section of Neoproterozoic Ophiolites from the Pan-African Belt, Eastern Desert, Egypt: Tectonomagmatic Evolution, Metamorphism, and Mineralization

Abstract

The Eastern Desert (ED) Neoproterozoic ophiolites are tectonically important elements of the Arabian–Nubian Shield. Although affected by various degrees of dismemberment, metamorphism, and alteration, almost all of the diagnostic Penrose-type ophiolite components can be found, namely, lower units of serpentinized peridotite tectonite and cumulate ultramafics and upper units of layered and isotropic gabbros, plagiogranites, sheeted dykes and pillow lavas. The contacts between the lower unit (mantle section) and the upper unit (crustal section) were originally magmatic, but in all cases are now disrupted by tectonism. The mantle sections of the ED ophiolites are exposed as folded thrust sheets bearing important and distinctive lithologies of serpentinized peridotites of harzburgite and dunite protoliths with occasional podiform chromitites. The ED ophiolites show a spatial and temporal association with suture zones that indicate fossil subduction zone locations. Multiple episodes of regional metamorphism mostly reached greenschist facies with less common amphibolite facies localities. CO₂-metasomatism resulted in the development of talc–carbonate, listvenite, magnesite, and other carbonate-bearing meta-ultramafic rocks. Geochemical data from the ED serpentinites, despite some confounding effects of hydration and alteration, resemble modern oceanic peridotites. The ED serpentinites show high LOI (≤20 wt%); Mg# mostly higher than 0.89; enrichment of Ni, Cr, and Co; depletion of Al₂O₃ and CaO; and nearly flat, depleted, and unfractionated chondrite-normalized REE patterns. The modal abundance of clinopyroxene is very low if it is present at all. Chromian spinel survived metamorphism and is widely used as the most reliable petrogenetic and geotectonic indicator in the ED ophiolite mantle sections. The high-Cr# (mostly ~0.7) and low-TiO₂ (mostly ≤ 0.1 wt%) characters of chromian spinel indicate a high degree of partial melt extraction (≥30%), which is commonly associated with fore-arc settings and equilibration with boninite-like or high-Mg tholeiite melts. Based on the general petrological characteristics, the ED ophiolitic chromitites are largely similar to Phanerozoic examples that have been attributed to melt–peridotite interaction and subsequent melt mixing in fore-arc settings. The comparison between the ED Neoproterozoic mantle peridotites and Phanerozoic equivalents indicates considerable similarity in tectonomagmatic processes and does not support any major changes in the geothermal regime of subduction zones on Earth since the Neoproterozoic era. The mantle sections of ED ophiolites are worthy targets for mining and exploration, hosting a variety of ores (chromite, gold, and iron/nickel laterites) and industrial minerals (talc, asbestos, and serpentine).

12.1 Introduction

Earth’s mantle is believed to consist of a heterogeneous assortment of ultramafic rocks, ranging in composition from fertile lherzolite to depleted harzburgite and ultimately, in some circumstances, to dunite (e.g., Furnes et al. 2014). Primitive, fertile mantle lithologies become depleted as a result of partial melting events in a variety of magmatic environments such as mid-ocean ridges and volcanic island arcs. The magmas resulting from such melting events are extracted and leave behind residual or depleted mantle lithologies (e.g., Jaques and Green 1980; Takahashi and Kushiro 1983; Winter 2001). Successive episodes of partial melting or reactions between peridotite and migrating melts can further deplete the residual mantle and alter the composition to highly depleted harzburgite or dunite. Note that dunite, though difficult to reach by simple melt extraction (Moores and Vine 1971), can be formed as a cumulate (e.g., Malpas 1978) or by replacive mechanisms (e.g., Quick 1981; Kelemen 1990; Kelemen et al. 1995; Allan and Dick 1996; Arai and Matsukage 1996). Areas where asthenospheric mantle reaches the lowest depths (and pressures), as well as areas affected by the addition of volatile components such as H₂O are expected to experience the largest degrees of partial melting (e.g., Kushiro 1969; Jaques and Green 1980) and so are favorable sites formation of dunite and chromitite (e.g., Quick 1981; Arai and Yurimoto 1995; Kelemen et al. 1995).

Mantle rocks can be studied in four main types of locality. Alpine-type or massif peridotites are found along continental margins and island arcs and in ancient orogenic belts as rootless lenticular bodies of serpentinized peridotite of various shapes and sizes (e.g., Coleman 1971; Jackson and Thayer 1972; Dick 1977; Sinton 1977; Arai 1980; Dilek and Newcomb 2003; and others). The Alpine-type massifs, with their extensive outcrop areas and often good exposure, provide valuable information about field relations between different mantle lithologies and give an estimate of the relative proportions of peridotitic rock species in the mantle. However, they have often experienced extensive lithospheric and surface processes that may obscure features of their original mantle history. Mantle xenoliths, by contrast, are often very fresh and directly sampled from depth. But they provide only isolated “snapshots” of a single vertical section beneath their eruption locality. Abyssal peridotites dredged directly from the seafloor, often in fracture zones along mid-ocean ridges, can be directly associated with particular mid-ocean ridge settings but are often highly altered. Intact ophiolites represent a fourth type of peridotite exposure. They have a well-known sequence of petrologic types with distinctive and easily recognizable mantle sections. When extensive tectonic dismemberment has removed them from their overlying crustal sections, as in most of the abundant examples in the Arabian–Nubian Shield (ANS), their outcrop patterns resemble those of Alpine-type peridotites. However, their characteristics allow in most cases confident assignment as the basal parts of incomplete ophiolite sequences. Each of these types of peridotite exposure can provide critical information about geochemical processes in the shallow mantle. Such processes include melt extraction, melt-rock reaction, and interaction with subduction-related fluids (e.g., Downes et al. 2001).

Study of mantle peridotites whose geodynamic and paleotectonic history are known has allowed the identification of diagnostic petrological characteristics that can, in turn, be used (with due caution) to infer the setting of ancient peridotites displaced from the context of their formation (e.g., Dick and Bullen 1984; Arai 1994a; Zhou et al. 1996; Ahmed 2013; Gahlan et al. 2015a; Gamal El Dien et al. 2016; and others). However, a range of different tectonomagmatic hypotheses have been proposed for the origin of the ED ophiolitic peridotites including (i) mid-ocean ridge (MOR) (e.g., Zimmer et al. 1995; Khalil 2007); (ii) back-arc basin (e.g., Ahmed et al. 2001; Farahat et al. 2004; Abd El-Rahman et al. 2009a; Farahat 2010); (iii) transitional back-arc basin-MORB (e.g., Abdel-Karim et al. 1996); and (iv) fore-arc (e.g., Azer and Stern 2007; Abd El-Rahman et al. 2009b; Ahmed 2013; Azer et al. 2013, 2019; Gahlan et al. 2015b).

Almost all the primary mantle minerals of the ED ophiolitic peridotites have been replaced during serpentinization or metamorphosis, except for relics of spinel and rarely of pyroxene or olivine. Hence, much of the reasoning leading to the assignment of tectonomagmatic settings for these rocks has been based on the chemistry of intact primary spinel (e.g., Dick and Bullen 1984; Arai 1992 1994a; Zhou et al. 1996). Depleted mantle residual to high degrees of partial melt extraction is most often found in the fore-arc of subduction zones and characterized by Cr-rich spinel (Cr# = molar Cr/[Cr + Al] > 0.6). On the other hand, MOR settings unrelated to subduction zones typically leave more fertile residues that have experienced lower degrees of partial melt extraction and contain more Al-rich spinel (Cr# < 0.6) (e.g., Dick and Bullen 1984; Arai 1994b; Zhou et al. 1996).

More information on the Neoproterozoic-age ophiolitic mantle peridotites is needed, particularly since the suggestion of Wynne-Edwards (1976) that there have been important changes in Earth’s geothermal regime since the upper Proterozoic. If issues related to severe serpentinization and metamorphism can be overcome, the mantle sections of the Neoproterozoic ophiolites of the Egyptian Eastern Desert (ED) are of great scientific interest for understanding such general questions of the geodynamic evolution of the mantle as well as, more locally, the tectonomagmatic processes beneath the ANS and along the Pan-African Belt during the Neoproterozoic era (e.g., Ahmed 2013; Gahlan et al. 2015b; Gamal El Dien et al. 2016; among others).

This contribution is an overview of the geology of the Neoproterozoic ophiolitic mantle peridotites exposed in the ED of Egypt. Field relations, petrography, mineral chemistry, and whole-rock geochemistry are integrated to define their general petrological characteristics and hence to constrain the geodynamic settings of the ophiolites. To examine the change in the geothermal regime, if any, from the Neoproterozoic to Phanerozoic era, we compare the Neoproterozoic ophiolitic peridotites of the ED to their Phanerozoic equivalents. Furthermore, given the striking metamorphism, alteration, and mineralization associated with the ED ophiolitic peridotites, we offer a discussion of types and grades of alteration and metamorphism on display and summarize their associated economic potential.

12.2 Geological Outline

The ANS is part of the East African Orogen, whose complex history includes records of the break-up of Rodinia at circa 900–800 Ma; and the evolution of numerous arc systems, oceanic plateaux, oceanic crust and sedimentary basins (e.g., Stern 1994; Stein and Goldstein 1996; Kusky et al. 2003; Kusky 2004); and obduction along convergent plate boundaries during the Pan-African orogeny (~750–650 Ma) (Stern 1994; Ali et al. 2010). The ANS in particular was assembled during the Neoproterozoic closure of the ~870–690 Ma Mozambique Ocean, and hence hosts a number of ophiolite-decorated sutures. In fact, the ANS may have the highest density of ophiolites of any Proterozoic terrane on Earth (e.g., Kusky et al. 2003; Kusky 2004; Stern et al. 2004). The occurrence of ophiolites decorating suture zones and their association with calc-alkaline rocks of island-arc affinity led several authors to frame their models of the Pan-African belts of the ANS in an Upper Proterozoic plate-tectonic context (e.g., Bakor et al. 1976; Garson and Shalaby 1976; Greenwoods et al. 1976, 1977, 1981; Shackleton et al. 1980).

Precambrian basement is exposed over an area of ~100,000 km² in the Red Sea Hills of the ED of Egypt (Fig. 12.1), and in the south Sinai Peninsula. Of this part of the ANS, the northern half of the East African Orogen, about 5.3% of the outcrop area is estimated to be ultramafic bodies (Dixon 1979; El Gaby et al. 1990; Hassan and Hashad 1990). The basement of the ED is comprised of four main tectonostratigraphic units. From bottom to top, these are (i) basal gneisses and migmatites, (ii) arc-type volcanic, volcano-sedimentary, plutonic, and ophiolite assemblages, (iii) the Ediacaran Hammamat and Dokhan supracrustal sequences, and (iv) syn- to late-tectonic granitoids that intrude all units (e.g., Ali et al. 2012; Hamimi et al. 2019). The ophiolites of the second main unit span an age range of about 200 Ma of ocean-floor magmatism (from 890 to 690 Ma) and up to 100 Ma (from 780 to 680 Ma) of terrane convergence and suturing (e.g., Stern et al. 2004; Ali et al. 2010). Some particular well-dated cases in Egypt, however, are concentrated in age span from ~730–750 Ma: the Gerf ophiolite at 741 ± 21 Ma (Kröner et al. 1992, Zimmer et al. 1995), the Ghadir ophiolite at 746 ± 19 Ma (Kröner et al. 1992), the Fawakhir ophiolite at 736.5 ± 1.5 Ma (Andersen et al. 2009), and the Allaqi ophiolite at ~730 Ma (Ali et al. 2010).

Fig. 12.1

Regional geological map showing the distribution of late Neoproterozoic ophiolitic rocks in the central and south Eastern Desert of Egypt (modified after Shackleton 1994)

The tectonostratigraphic column of the Egyptian ophiolites can, if assembled, display the complete sequence of the Penrose Conference ophiolite model (Anonymous 1972), from mantle section upward to the overlying mafic plutonic sequence and then mafic volcanic rocks (Fig. 12.2). However, all ophiolites of the Egyptian ED have been variably dismembered by later tectonism. The isolated serpentinite masses in the ED were first defined as ophiolites by Rittmann (1958), although the term ophiolite in the sense of the contemporary definition was first applied to the ED cases by Bakor et al. (1976), Garson and Shalaby (1976), and Neary et al. (1976). Based on the field experiences of the authors, personal communications with other Egyptian geologists, and the published record, we have compiled a list of the best preserved and most complete ophiolites of the ED in Table 12.1.

Fig. 12.2

Schematic tectonostratigraphic column of the Egyptian ophiolites (adapted from Gahlan et al. 2015b)

Table 12.1 The best preserved Pan-African ophiolites in the Eastern Desert of Egypt

Let us consider the mantle section of the Gerf ophiolite as a typical example of the appearance of these bodies. The mantle section is exposed in massive ridges and sheet-like bodies of serpentinized harzburgite, dunite, and rarely lherzolite. It is fault-bounded and well preserved in the troughs of major synforms (Fig. 12.3). The contact against country rocks is structural, defined by thrust faults verging to the W and SW (Fig. 12.4), suggesting orogenic stresses acting generally from the E and NE during the collision of East and West Gondwana (e.g., Kröner et al. 1987; Stern et al. 1990; Shackleton 1994; Nasr et al. 1996; Abdelsalam et al. 2003; Abd El-Rahman et al. 2009a; and references therein). Any metamorphic sole that may have developed during obduction has been concealed by the overthrust mantle section, and no contact metamorphism is observed toward the footwall. The mantle section, in places, was affected by the intrusion of syn- to late-orogenic calc-alkaline granite plutons, with narrow thermal contact metamorphic aureoles at their contacts (e.g., Khalil and Azer 2007; Gahlan and Arai 2009; Ahmed et al. 2012a).

Fig. 12.3

Mosaic of field photos demonstrating massive serpentinites being preserved within the trough of a major synform at Wadi Hasium in the Gerf ophiolite nappe. Note that the original thrust contact between the low-grade metasediments (M) and the structurally overlying Gerf serpentinites (S) is folded and the eastern limb dips toward the west by 70° (270/70°) (adapted from Gahlan 2006)

Fig. 12.4

The Gerf serpentinites thrust over low-grade metasediments, both intruded by the Shinai syn- to late-orogenic granite, with talc–carbonates demarcating the thrust contact (Wadi Shinai, Gerf ophiolite nappe; adapted from Gahlan 2006)

Serpentinized harzburgite, the dominant mantle lithology, forms a screen that incorporates less abundant serpentinized dunite veins, layers, and lenses. Tectonic fabrics, foliations, and lineations are observed in the serpentinized harzburgite (Fig. 12.3). Toward the basal thrust and along faults and shear zones, the serpentinized peridotites are transformed into schistose serpentinites and talc–carbonates. The latter can be seen from a far distance as creamy-colored streaks cutting across the massive serpentinites. Concordant to sub-concordant small-scale chromitite pods, often observed in the uppermost parts of mantle sections (e.g., Ahmed et al. 2001; Gahlan et al. 2015b), are here enveloped by dunite that is, in turn, enclosed in the harzburgite screen (Fig. 12.2).

Large masses of serpentinized dunite, concordantly interlayered with serpentinized ultramafic cumulates (wehrlite and pyroxenite) and gabbro sills often define the uppermost part of ophiolite mantle sections (e.g., Ras Salatit ophiolite, Gahlan et al. 2012). This interval is called the Moho Transition Zone (MTZ) (Fig. 12.2). An interval of closely spaced gabbro sills often marks the uppermost part of the MTZ and may grade into the plutonic crustal section (Fig. 12.2). In the Gerf exposure, this initially magmatic contact between lower and upper units is now a structural contact due to severe tectonic disruption (Fig. 12.4).

Alteration of ophiolitic ultramafic rocks in the ED has led to the development of talc–carbonate, listvenite, magnesite, and (less commonly) rodingite. Talc–carbonate rocks are found both in the volcanic and mantle sections of ED ophiolites and so are divided metavolcanic-derived talc–carbonate and ultramafic-derived talc–carbonate. The majority of talc occurrences in Egypt are ultramafic-derived and may be found as separate bodies or associated with serpentinites. These rocks are characterized by their conspicuous buff/creamy color and cavernous nature. The cavernous appearance (Fig. 12.5) and porosity have been attributed to dissolution of the carbonate at or near the exposure surface by carbonic acid in meteoric water. They are generally rather uniform in mineralogical composition but vary appreciably in fabric, including massive, schistose, and gneissose varieties. The talc-rich rocks are commonly soft but they become progressively harder with increasing iron content or abundance of carbonate minerals. A few nodules and irregular pockets of magnesite and dolomite with quartz are observed in the talc-rich rocks. Some serpentinite masses are observed within the talc–carbonates (Fig. 12.6).

Fig. 12.5

Talc–carbonates at Barramiya showing their conspicuous buff/cream color and cavernous appearance

Fig. 12.6

Serpentinite mass within talc–carbonates at Barramiya

Listvenite bodies of various shapes and sizes are developed along shear zones. They express positive geomorphic relief relative to the surrounding rocks due to their resistance to arid-climate weathering. In poorly exposed areas, the presence of upstanding ridges of listvenite may be the only evidence for underlying altered ultramafic rocks. They mainly form dyke-like bodies or lenses hundreds to thousands of meters long (Fig. 12.7). Structural elements in the sheared listvenite rocks are generally conformable to the main plano-linear fabric of the host serpentinite. They are pink to reddish-brown in color due to iron-oxide staining. A few of the listvenite outcrops show porous textures due to supergene oxidative weathering. Locally, listvenites may be brecciated and fractured, with fractures being filled by carbonate veinlets and fine quartz ribbons.

Fig. 12.7

Listvenite dyke-like body developed along shear zones, Atud area

Magnesite forms veins, stockworks, and massive bodies of snow-white color in the serpentinite country rocks (Figs. 12.8 and 12.9) and along regional faults cutting the ultramafic rocks. The boundaries of the magnesite bodies are knife sharp but irregular. The magnesite veins are branched and have variable lengths and a range of widths from a few centimeters up to 1 m. Some brecciated serpentinites are present within the massive magnesite.

Fig. 12.8

Snow-white deposits of massive magnesite at Um Khariga

Fig. 12.9

Branched veins and stockwork of magnesite at Gabal Ghadir

Rodingite appears in the ED ophiolites within serpentinized ultramafics and ophiolitic mélange, but these occurrences are only sparsely documented in the literature (Takla et al. 1992; Abdel-Karim 2000; Surour 2019). Rodingite is distinctive in outcrop; its white color with black spots rock provides a conspicuous contrast with dark-colored ultramafic host rocks. In the Gerf case, for example, rodingite appears rather homogeneous in the field, but two types of rodingite are recognized in ED ophiolites: Type I rodingite forms thin dykes in serpentinized peridotite (Fig. 12.10), whereas Type II rodingite forms blocks and irregular lenses in ophiolitic mélange (Azer, unpublished data). Narrow chloritite rims (blackwall) can be found between Type I rodingite and serpentinite, representing a transitional reaction zone between the two rock types. Some small serpentinized ultramafic xenoliths occur within Type II rodingite.

Fig. 12.10

Rodingite dyke in serpentinized peridotites, Um Rashid area.

12.3 Structural Setting

In many areas in the ED (e.g., Beitan, Hafafit, Barramiya, El-Shalul, Meatiq, and Sibai), well-developed Pan-African thrusting can be observed separating hanging wall supracrustal unit (ophiolites and ophiolitic mélange) from infracrustal gneisses, migmatites, sheared granites, and remobilized equivalents in its footwall. Several lines of evidence support the interpretation of these contacts as thrust faults, including the contrast in metamorphic grade between supracrustal and infracrustal rocks, narrow mylonite zones, intensity of shearing and tightness of foliation, and the cataclastic nature of both litho-tectonic units close to the thrust contact. The supracrustal unit consists of dismembered blocks and masses of serpentinite, metagabbro, and metavolcanics, embedded within a matrix of low-grade metasedimentary rocks including met siltstones, slates, phyllites, and schists (Fig. 12.11). The serpentinite blocks vary in size, from small pebbles to large blocks, and may be enclosed either in a matrix of metasediments or within highly sheared serpentinite, with block boundaries usually marked by cataclasis and/or mylonitization.

Fig. 12.11

Close-up view showing shearing-related late chevron folding within tremolite-actinolite schist in the Hafafit Core Complex. The fold axis plunges NW (after Hamimi et al. 2018; Abo Soliman 2019)

Large mountainous nappes are encountered in Gerf, Muqsim, Shilman, Abu Dahr, Um Taghar, Barramiya, Shalul, and in many other areas. The Gerf ophiolitic nappe, for example, crops out immediately north of the N-to-NNE trending Hamisana Shear Zone. It is the largest nappe in the ANS, some 30–40 km in diameter, and preserves a complete N-MORB ophiolite sequence (Zimmer et al. 1995). A single zircon Pb/Pb age of 726 ± 40 Ma and a whole-rock Rb/Sr age of 551 ± 28 Ma are given by Stern et al. (1989) for a post-nappe granodioritic intrusion, whereas gabbros and basalts associated with the Gerf ophiolite give Sm/Nd whole-rock ages of 720 ± 9 Ma and 758 ± 34 Ma, respectively (Zimmer et al.,1995). Kröner et al. (1992) obtained a single zircon Pb/Pb age of 741 ± 21 Ma for layered gabbro from the Gerf ophiolite. Such ages led Zimmer et al. (1995) to conclude that the Gerf ophiolite was emplaced due to collision of island-arc complexes between 600 and 700 Ma.

The ophiolitic nappes occasionally contain variably sized blocks of amphibolite (Fig. 12.12), metagabbro, metavolcanics, and chromitite lenses. In Abu Dahr and Sefein, chromitite lenses have been systematically oriented by the effects of deformation. The ophiolitic nappes are frequently traversed by NW- (and NE-) oriented Najd-related transpressional shear zones. In many cases, shearing led to the formation of well-developed shear zone-related folds. Kinematic indicators with monoclinic symmetry and slicken lines (Fig. 12.13) reflect the overall kinematics along these Najd shears. In addition, piggyback thrusts and thrust-related structures imbricate thrust stacks (Fig. 12.14), and thrust duplexes are recorded elsewhere in the ophiolitic nappes. Thrusting propagated according to the “footwall-nucleating-footwall-vergent rule,” where earlier hanging walls are carried forward in a piggyback manner, and newly formed thrusts grow in the footwalls of the older thrusts. In places, out-of-sequence thrusts are documented. Overprinting relations demonstrate that, in most cases, thrust-related folds are older than shear zone-related folds, as documented in many deformed belts in the ED including the Mubarak–Baramiya belt, the Um Nar–Gabal Elhadid Belt, and the Wadi Khuda Belt (Hamimi et al. 2019). Thrust planes are often marked by magnesite, talc, chrysotile, and tremolite asbestos. Magnesite forms veins, stockworks, and massive bodies of snow-white color that are sometimes folded (Fig. 12.15). Far from the thrust planes, larger serpentinite blocks feature massive cores and sheared outer parts. Massive serpentinite is dark gray in color and slightly deformed. Increasing shearing intensity and tectonic transport lead to brecciated serpentinite. Highly sheared serpentinite is usually yellow to pale green. The smaller blocks and the peripheries of the large blocks show widespread shearing and milling down of the serpentinite, resulting in the formation of schistose tectonic serpentinite matrix.

Fig. 12.12

Amphibolite block in highly sheared meta-ultramafic rocks, the northern part of Nugrus shear zone

Fig. 12.13

Slickenlines encountered along a NW-oriented shear zone, west of Meatiq dome

Fig. 12.14

Imbricated thrust stacks in highly sheared meta-ultramafic rocks near the Kordman Gold Mine

Fig. 12.15

Folded magnesite in the vicinity of Gabal Zabar

In the extreme southern part of the ED, ophiolites form elongated blocks and nappes decorating the Allaqi–Heiani high-strain zone, interpreted to represent the western continuation of the conspicuous Allaqi–Onib–Sol Hamed–Yanbu arc–arc suture. The E–W extension of this ophiolite-decorated deformation suture defines the location and orientation of a fossil subduction zone between the Midyan-ED Terrane to the north and Hijaz–Gebeit–Gabgaba Terrane to the south. It is regarded by many workers (e.g., Abelsalam et al. 2003; Hamimi et al. 2019; Fowler and Hamimi 2020; Hamimi and Abd El-Wahed 2020) as a key to understanding the Neoproterozoic evolution not only of the Egyptian Nubian Shield but also the ANS as a whole.

Tectonically speaking, the rates and directions of transportation of the ophiolitic nappes have been the subject matter of detailed discussions by many workers (e.g., Greiling 1985, 1997; Abdelkhalek et al. 1992; Hamimi 1996; Abdelsalam et al. 2003; Hamimi et al. 2019; Fowler and Hamimi 2020). In these papers and references therein, one finds two proposed transport directions: W-to-WSW and N-to-NNW. The W-to-WSW transport direction is most probably a consequence of the final assembly and accretion of the whole ANS to the Saharan Metacraton (Abdelsalam et al. 2011) which was concurrent with the assembly of eastern and western Gondwana during the late Cryogenian–Ediacaran (650–542 Ma). The N-to-NNW transport direction, on the other hand, is associated with the N-directed tectonic escape of the ANS and is best preserved as N-to-NNW trending stretching lineations in ophiolitic nappes and huge blocks. The rates of transport of the ophiolitic nappes in the ED varied considerably from one deformed belt to another, depending on many parameters, including nappe size, shearing rate, and the rheology of the enclosing lithologies (Hamimi et al. 2019).

12.4 Petrography

Petrography of the ED ophiolitic mantle lithologies will be described generally in the light of their field observations. Despite extensive replacement by serpentine minerals, an experienced petrographer identifies primary mantle mineral modes with the aid of relict textures, supported by chromian spinel chemistry and morphology (e.g., Arai 1980; Liipo et al. 1995; Matsumoto and Arai 2001). Harzburgite, dunitic harzburgite, and dunite are the predominant protoliths. Although serpentinization and alteration of the ED mantle peridotite rocks are pervasive, each process is variable in intensity. Considerable local differences of H₂O and CO₂-activity during metamorphism of ED mantle sections led to the development of various categories of metaperidotite, including lizardite serpentinite, carbonated metaperidotite, magnesite–antigorite serpentinite, anthophyllite-bearing serpentinite, and talc–carbonate rocks (e.g., Gahlan and Arai 2009; Ahmed et al. 2012a; Gahlan et al. 2015b among others). The term serpentinite is applied for a rock composed of ≥50% serpentine-group minerals (e.g., Coleman 1971), and likewise carbonate-bearing rocks as rocks containing ≥50% carbonates. Widespread oxidative low-temperature alteration lends the pervasive reddish-creamy color to many carbonate-bearing outcrops (Fig. 12.16). Along shear zones and around igneous intrusions, the serpentinites are transformed into talc–carbonates and carbonated metamorphic peridotites, where the original textures are partly or completely obliterated by recrystallization (e.g., Gahlan and Arai 2009). Moreover, intensive shearing resulted in crudely foliated fine clastic rocks.

Fig. 12.16

Peridotite pervasively altered to reddish-cream carbonates at Barramiya

Lizardite serpentinites of harzburgite parentage are the most common rock type in the ED ophiolite mantle sections and represent the lowest degree of metamorphism. They are usually massive, fine-grained, and dark grayish-green in color with green patches and yellowish-green weathered surfaces (Fig. 12.17). Petrographically, they consist mainly of serpentine minerals (lizardite/chrysotile), variable amounts of carbonate minerals (magnesite and dolomite), subordinate relics of primary mantle minerals (olivine, pyroxene, and chromian spinel), and accessory sulfides and magnetite. They are dominated by pseudomorphic microtextures (e.g., O’Hanley 1996).

Fig. 12.17

Yellowish-green weathered surfaces of lizardite serpentinites of harzburgite parentage at Gabal Ghadir

Generally, harzburgite protoliths had almost exclusively protogranular textures except those close to shear zones and thrust contacts. Even when completely serpentinized, it is possible to recognize pseudomorphs after primary orthopyroxene (bastite texture, Fig. 12.18a) and olivine (mesh texture, Fig. 12.18b). Bastite is typically low in modal abundance (≤20 vol.%); the subhedral shapes and uniform texture indicate one precursor pyroxene, probably orthopyroxene (Fig. 12.18a). Most bastites are composed of lizardite with lesser chrysotile (e.g., Whittaker and Zussman 1956; Page 1967); we have never seen antigorite bastite, despite some published descriptions (Hess et al. 1952; Wicks and Whittaker 1977). Chromian spinel forms deep reddish-brown, anhedral, and embayed crystals with submetallic luster in reflected light. Alteration products with higher reflectivity and sharp optical boundaries with unaltered parts are developed along rims and cracks of chromian spinel (Fig. 12.18c), often in order from ferritchromite through Cr–magnetite to magnetite. The degree of spinel alteration is commonly higher in serpentinized harzburgite than in dunite. Chromian spinel commonly includes Al–serpentine, chlorite, olivine, dolomite, calcite, or sulfides.

Fig. 12.18

Photomicrographs showing petrographic textures: (a) bastite texture within serpentine minerals; (b) mesh texture within serpentine minerals; (c) chromian spinel crystal altered along the margins and cracks to ferritchromite; (d) fresh relics of olivine in serpentinized dunite; (e) fresh relics of orthopyroxene in serpentinized harzburgite; (f) antigorite showing feather, rosette, and flame textures with carbonates; (g) chrysotile veinlets cutting antigorite; and (h) euhedral chromian spinel crystals in serpentinized peridotite. All images in cross-polarized transmitted light except (a) in plane-polarized transmitted light, (c) in reflected lighted, and (h) is backscatter image

Antigorite serpentinites are generally massive and fine-grained. They occur as dark greenish-gray lenticular layers and lenses within serpentinized harzburgite. Occasionally, they are dissected by parallel sets of very small-scale oriented shear zones and fractures filled with carbonates (mostly magnesite). Petrographically, they consist mainly of antigorite with or without fresh relics of primary silicates, magnesite (rarely dolomite), and subordinate amounts of chromian spinel, magnetite, and sulfides (mainly pentlandite). Relics of primary minerals in antigorite serpentinite from the Gerf ophiolite, e.g., include dark brown patches of anhedral, cracked olivine (Fig. 12.18d), and pyroxene (Fig. 12.18e) set in a dense fine-grained groundmass of serpentine (Gahlan et al. 2015b). Antigorite flakes are occasionally replaced by (or included in) carbonate minerals, principally magnesite. Larger antigorite crystals host magnetite dust or fine granules. The process of recrystallization to antigorite was accompanied by the development of non-pseudomorphic interpenetrating fabrics (e.g., O’Hanley 1996), including rosette or knitted structures, flame texture of Francis (1956), and thorn texture of Green (1961) (Fig. 12.18f). Olivine, if any, is found as relics set in a dense groundmass of antigorite. Chrysotile develops as veinlets (Fig. 12.18g) or along relict olivine fractures and cracks; the margins of chrysotile veins are embayed by antigorite. Deep reddish-brown chromian spinel forms euhedral to rounded crystals, which is a characteristic crystal form of spinel inherited from dunite protoliths (Fig. 12.18h). The degree of spinel alteration is commonly lower than that in harzburgite. Alteration to Cr–magnetite and magnetite, with less developed ferritchromite, is found only along grain boundaries and cracks. Sulfides, chlorite, magnesite, dolomite, and Cr–tremolite are commonly included in chromian spinel. Magnetite forms anhedral equant grains, randomly distributed throughout the rocks or aligned in veinlets along cracks. Magnetite is less abundant in serpentinized dunite than in harzburgite.

Both massive and disseminated chromitites are found in the mantle sections of ED ophiolites. Massive chromitite (Fig. 12.19a) is composed mainly of coarse, equidimensional, euhedral to anhedral crystals of chromian spinel (>92 vol.%) with some interstitial serpentine minerals and rare olivine. Chrome spinel is commonly unaltered or may show only slight alteration to ferritchromite along grain margins and cracks. Very small olivine inclusions are observed within the chromian spinel. Cumulate, chain structures, and banding are the most common textures of the massive chromitite, typical of magmatic crystallization (Pal and Mitra 2004). Disseminated chromitite (Fig. 12.19b) consists of chromian spinel (40–70 vol.%), intergranular serpentine minerals and rare olivine, chlorite, and opaques. Compared to the massive chromitite, most chromian spinel grains in disseminated chromitite are more euhedral, smaller, and more altered, with highly porous rims of ferritchromite. The greater intensity of alteration in disseminated chromitites can be attributed to easier subsolidus elemental redistribution with neighboring silicate phases, which are nearly absent in massive chromitite.

Fig. 12.19

Photomicrographs of alteration features: (a) massive chromitite with minor serpentine along grain boundaries and cracks, (b) disseminated chromitite, (c) fine aggregates of talc, (d) listvenite with fuchsite, (e) cracked chromian spinel crystals within massive magnesite, and (f) anthophyllite rock. All images in plane-polarized transmitted light except (c, f) are in cross-polarized transmitted light

The ED ophiolite mantle sections also feature a number of distinctive alteration products, including talc–carbonates, listvenite, and magnesite with less common anthophyllite and rodingite rocks. Metasomatic alteration of serpentinite to talc–carbonates occurs to variable extents. The talc–carbonate rocks are fine-grained with brownish-yellow to reddish-brown color. They are composed essentially of talc (>75 vol.%) with subordinate serpentine, carbonate, and opaques (Fig. 12.19c). Talc occurs as fine aggregates or as platy or micaceous grains. Carbonates, principally magnesite with minor dolomite and occasional calcite, are irregularly distributed and form coarse-grained aggregates. Opaque minerals include magnetite, chromian spinel, and sulfides.

Listvenite (also spelled “listwaenite” and “listwanite”) is an unusual rock type with a distinctive mineralogy of Cr-rich muscovite (fuchsite), quartz, and carbonates. It represents the end product of carbonatization, potassic alteration, and silicification (Halls and Zhao 1995; Azer 2013; Gahlan et al. 2018). Petrographically, Azer (2013) and Gahlan et al. (2018) distinguished listvenite into two types, silica-rich and carbonate-rich. The presence of fuchsite in the silica-rich variety indicates that it is typical listvenite, while the absence of fuchsite in the carbonate-rich variety suggests that “listvenite-like rock” is a more precise name for it. Most likely these two types of listvenite are formed by alteration under different conditions or in fluids of different compositions, rather than trapping progressive states along a single alteration path. In general, listvenites are composed of carbonates (magnesite and breunnerite with minor calcite), quartz, Fe–Ti oxides, serpentine, chromian spinel, and chlorite. The Si-rich variety also features the characteristic fuchsite as an accessory mineral in the form of flakes, fine disseminated crystals, and thin bands with a perfect cleavage in one direction (Fig. 12.19d).

Magnesite ores in the ED occur as veins and massive bodies associated with serpentinite. Vein magnesite is pure magnesite and predominantly cryptocrystalline in texture, with sharp contacts against enclosing serpentinites. Rare angular fragments of host rocks are observed near some vein margins but there are no other silicate minerals in the veins. Massive magnesite consists essentially of magnesite with minor dolomite and calcite. Rare chromian spinel crystals similar to those in the host rocks occur within the massive magnesite (Fig. 12.19e). Angular to sub-angular fragments of serpentinite and rare quartz veinlets are observed within massive magnesite pods.

Anthophyllite rock was formed due to the alteration of harzburgite through the introduction of SiO₂-rich fluids derived from the country rocks. It was formed at temperatures around 500 °C (Azer, unpublished data). Anthophyllite rock is composed mainly of fibrous anthophyllite with subordinate amphibole and opaque minerals (Fig. 12.19f).

12.5 Geochemistry

Most published chemical analyses from the mantle sections of ED ophiolites are from serpentinite samples (e.g., Zimmer et al. 1995; Abd El-Rahman et al. 2009a, b; Azer et al. 2013; Khalil et al. 2014; Obeid et al. 2016; Boskabadi et al. 2017; Abdel-Karim et al. 2018; Gahlan et al. 2018; among others). All analyzed serpentine samples contain abundant water and carbonate, with high LOI (typically greater than 10% and less than 20%). They have Mg# (100*Mg/[Mg + Fe^T] on a molar basis) greater than from 89, as do modern oceanic peridotites (Bonatti and Michael 1989). The typical Mg# values and very low K₂O and Na₂O contents indicate limited elemental mobility during serpentinization. Also, the very low concentrations of CaO in carbonate-free serpentinites are representative of metamorphic ophiolitic peridotites (Coleman 1977), as opposed to ultramafic cumulates.

Due to the extensive replacement of primary minerals by serpentine, estimates of the modal percentages of primary minerals are imprecise and classification by a modal scheme such as Streckeisen (1976) can only be approximate. Some authors have instead applied a classification using normative mineralogy based on whole-rock chemistry, neglecting any changes to anhydrous composition that may have accompanied serpentinization (e.g., Azer and Khalil 2005; Gahlan et al. 2015a; Khalil et al. 2014; Obeid et al. 2016; Azer et al. 2019). This places the majority of serpentinized ED ultramafic samples in the harzburgite and dunite fields, with rare lherzolite (Fig. 12.20a), in agreement with petrographic and field assessments.

Fig. 12.20

a Nomenclature of ED serpentinized ultramafic rocks based on Ol–Opx–Cpx normative composition, compared to field and petrographic assignment (after Coleman 1977) and b Al₂O₃ versus CaO diagram for ED serpentinized ultramafic rocks compared to fields of Ishii et al. (1992)

Geochemical data indicate that addition or subtraction of elements other than water and perhaps silica was very limited for massive serpentinites (i.e., away from shear zones and faults), encouraging us to compare these compositions with those of peridotites from modern tectonic settings. The very low abundance of Al₂O₃ in the ED serpentinites (Azer and Stern 2007; Abd El-Rahman et al. 2009a, b; Gahlan et al. 2015b; Obeid et al. 2016; Abdel-Karim and El-Shafei 2018) resembles “oceanic trench” peridotites as defined by Bonatti and Michael (1989). Similarly, the low mean CaO (<1%) suggests that protoliths were poor in clinopyroxene and affinities with much depleted oceanic peridotites (Bonatti and Michael 1989). Moreover, on the Al₂O₃ versus CaO diagram, the Egyptian serpentinites plot with harzburgites recovered from modern intra-oceanic fore-arcs (Fig. 12.20b).

As expected for peridotites, the trace element contents of the Egyptian serpentinites are highly variable, but they are uniformly depleted in most trace elements with unfractionated flat REE pattern, and enriched in the compatible elements Cr, Ni, and Co.

12.6 Discussion

12.6.1 Protolith and Geodynamic Setting

The ophiolitic rocks of the ANS have long been the subject of research because they represent important elements for reconstructing the geodynamic evolution of the Pan-African belt. In the northern ANS, ophiolites on both sides of the Red Sea are generally interpreted to have formed in suprasubduction zone (SSZ) tectonic settings (e.g., Nassief et al. 1984; Pallister et al., 1989; El Sayed et al. 1999; Ahmed et al. 2001; Farahat et al. 2004; Stern et al. 2004; Azer and Stern 2007; Abd El-Rahman et al. 2009b). Seafloor spreading is necessary to form SSZ ophiolites, whether in the fore-arc during the infant arc stage of subduction initiation or in a back-arc basin (Pearce 2003; Stern 2004). Resolving which of these environments is represented by a particular SSZ ophiolite is important because fore-arc ophiolites mark the formation of new subduction zones (Stern 2004) in episodes often associated with major plate reorganizations. In contrast, back-arc basins can form at any time in the evolution of a convergent plate margin.

Most assessments of tectonic setting for ED ophiolites have focused on the trace element composition of basic rocks (lavas and gabbros) and have rarely considered the abundant mantle sections. However, interpretation of tectonic setting for Neoproterozoic ophiolitic rocks on the basis of metavolcanic and metagabbro samples encounters difficulties due to the effects of fractional crystallization and alteration. Even when these problems are minimized, it can be very difficult to distinguish the chemical compositions of fore-arc and back-arc lavas (Azer and Stern 2007). The geochemistry of lavas of the ED ophiolites appear to be transitional between island-arc basalt and MORB, which has led a number of authors to back-arc environments for Egyptian ophiolites (e.g., El Sayed et al. 1999; Farahat et al. 2004; Abd El-Rahman et al. 2009a). However, the transitional character of the ophiolitic lavas is simply due to their hydrous nature and the effects of added slab-derived components (Azer and Stern 2007), which is not restricted to the back-arc. Building of the relatively recent discovery and exploration of fore-arc spreading during subduction initiation (Shervais et al. 2004; Stern 2004), a number of recent papers have instead argued for a fore-arc setting for the Egyptian ophiolites (Azer and Stern 2007; Khalil and Azer 2007; Gahlan et al. 2015a 2018; Azer et al. 2019; among others).

Although controversy over the detailed tectonic environment of the ED ophiolites, the case for fore-arc affinity is strong. Phanerozoic boninites appear to be restricted to fore-arc settings (e.g., Murton 1989; Johnson and Fryer, 1990; Bédard 1999; Beccaluva et al. 2004). Magmas with boninitic affinities have been reported from the ANS (Wolde et al. 1993; Yibas et al. 2003; Katz et al. 2004; Teklay 2006), and in particular such affinity has been found in mafic members of some ED ophiolites (e.g., El Sayed et al. 1999; Abdel Aal et al. 2003; Saleh 2006).

In order to use evidence from the mantle sequences of the ED ophiolites to gain further insight into tectonic affinity, one must first consider possible artifacts. Serpentinization has obviously affected the ultramafic units of the ED ophiolites, but in many cases serpentinization may be an essentially isochemical process (Coleman and Keith 1971; Donaldson 1981; Shiga 1983, 1987), except for Ca and mobile incompatible trace elements have been also leached from the peridotites during serpentinization. Ni, Cr, Co, and V, in particular, are compatible and relatively “immobile” during alteration (e.g., Hébert et al. 1990; Pereira et al. 2003). In the ED cases, bulk serpentinites show uniformly high Mg# and enrichment in Ni, Cr, and Co. Together the high Cr# (up to 0.8) of the intact spinel, these features all suggest high degrees of partial melt extraction (e.g., Gülacar and Delaloye 1976; Dick and Bullen 1984; Arai 1994a; Ishiwatari et al. 2003).

The marked depletion observed in incompatible elements, as well as the nearly flat, depleted, and unfractionated REE chondrite-normalized pattern of the ED serpentinized peridotites indicates highly refractory peridotite protolith (harzburgite and dunite), high degrees of partial melting and high melt/rock ratio (e.g., Roberts 1992; Bodinier and Godard 2003; Gahlan et al. 2015b; Obeid et al. 2016; among others), though it may also reflect later modifications. Moreover, the parental melt of the ED peridotites is suggested to be little fractionated by plagioclase crystallization, which is supported by the almost nil Eu* negative anomaly (e.g., Drouin et al. 2009; Obeid et al. 2016). The behavior of fluid-mobile elements, Au, As, S, LILE, and Pb in the ED serpentinites, could be attributed to alteration and/or the action of fluids produced by dehydration of subducted slab (e.g., Gahlan et al. 2015b; Boskabadi et al. 2017).

In some places, metamorphism, alteration, and deformation have been severe enough to obliterate primary textures and mineralogy. However, in most of the massive serpentinites, one can still identify primary mantle lithologies with the aid of relict textures (e.g., bastite and mesh serpentinite) and chromian spinel chemistry and morphology (e.g., Arai 1980; Liipo et al. 1995; Matsumoto and Arai 2001). The relict protogranular texture and vermicular shape of chromian spinel in many localities may indicate harzburgites that did not experience significant post-magmatic deformation.

Spinels are the only igneous minerals that retain most of their original chemistry in the serpentinized peridotites. Even in completely serpentinized ultramafic rocks containing no relics of primary silicate minerals, the chemical composition of unaltered accessory chromite has been widely recognized as a potentially important petrogenetic indicator (e.g., Barnes and Roeder 2001; Sobolev and Logvinova 2005; Arif and Jan 2006). Spinel from MOR and back-arc basin peridotites generally have Cr# <50 (Barnes and Roeder 2001; Ohara et al. 2002), whereas spinel in fore-arc peridotites generally has higher Cr# (up to 80) and spinel from boninites typically has Cr# of 70–90 (Ohara and Ishii 1998; Stern et al. 2004). The Cr# of spinel in ED ophiolitic peridotites is mostly >60 (e.g., Khalil and Azer 2007; Khalil et al. 2014; Gahlan et al. 2015b; Azer et al. 2019) and similar to those of modern fore-arc peridotites. The negative correlation between Cr# and Mg# of chromian spinel (Fig. 12.21a; see also Fig. 12.8 in the companion paper by Azer and Asimow in this volume) may reflect variable partition coefficients for Mg and Fe between chromian spinel and olivine or silicates (Irvine 1965; Dick and Bullen 1984).

Fig. 12.21

a Variation of TiO₂ (wt%) versus Cr# of chromian spinel in ED chromitite, dunite, and harzburgite. Fields of MORB and boninite are after Dick and Bullen (1984) and Arai (1992). Fields of depleted and highly depleted peridotites are after Jan and Windley (1990). Field of fore-arc peridotite (dashed) is after Ohara and Ishii (1998). b Comparison, in terms of spinel composition, between the Neoproterozoic ED peridotites and analogous Cenozoic ophiolite fields (Bloomer et al. 1995). The shaded field of the Neoproterozoic ANS peridotites is after Stern et al. (2004). Note that the vast majority of the ED ophiolitic peridotites and the associated chromitites plot in the field of ANS peridotite

The high Cr# of chromian spinel from the ED peridotites is combined with low TiO₂ contents, mostly <0.3 wt% (Fig. 12.21b). This feature is also consistent with high degrees of near-fractional partial melting and melt extraction (Mysen and Kushiro 1977), as expected for the mantle wedge in SSZ (fore-arc) settings (e.g., Dick and Bullen 1984; Bonatti and Michael 1989; Arai 1992, 1994a). Ohara and Ishii (1998) found similar accessory chromian spinel with high Cr#, up to 0.8, in fore-arc peridotites from the Mariana Trench. Thus, the spinel chemistry in the mantle sections Neoproterozoic ED ophiolites strongly supports the inferences previously drawn from their volcanic sections that they are fragments of oceanic lithosphere that experienced boninitic or high-Mg tholeiitic magmatism in an SSZ setting (e.g., Pearce et al. 1984; Arai 1992, 1994b).

The fresh relics of primary olivine found in most peridotites, even when strongly serpentinized, can provide further insight into the tectonic setting for the protoliths (Parkinson and Pearce 1998; Pearce et al. 2000; Coish and Gardner 2004). Relict olivine in Egyptian ED serpentinites is Mg-rich, with Fo content >88 (Khudeir 1995; Khalil and Azer 2007; Gahlan et al. 2015b; Khalil et al. 2014; Obeid et al. 2016; Gahlan et al. 2018, Azer et al. 2019), much like relict olivine analyzed in other ANS ophiolites (e.g., Ledru and Auge 1984; Nassief et al. 1984; Stern et al. 2004). In this regard, they resemble olivine in fore-arc peridotites (Arai 1994b), interpreted to be residues after extensive melt extraction. Compositions of coexisting olivine and spinel (Fig. 12.22a) in the ED ophiolitic serpentinites further supports a fore-arc setting for these ophiolites (e.g., Khalil and Azer 2007; Khalil et al. 2014; Gahlan et al. 2015b; Obeid et al. 2016; Gahlan et al. 2018, Azer et al 2019).

Fig. 12.22

a Cr# of spinel versus Fo content of coexisting olivine from serpentinized ED peridotites (Arai 1992). PM: Primitive mantle, OSMA: Olivine–spinel mantle array (Arai 1994a), b TiO₂–Na₂O–SiO₂/100 diagram for fresh relics of clinopyroxene in serpentinized ED peridotite (Beccaluva et al. 1989). WOPB = within-ocean plate basalts; MORB = mid-ocean ridge basalts; IAT = island-arc tholeiites; BON + BA-A = boninites + basaltic andesites and andesites from intra-oceanic fore-arcs

The low Al₂O₃ and high Mg# of fresh relics of orthopyroxene in serpentinized peridotites from the ED ophiolites are also consistent with those found in highly depleted fore-arc peridotites (e.g., Ishii et al. 1992; Bonatti et al. 1993). Finally, despite the scarcity of clinopyroxene in the highly depleted ED peridotites, exceedingly rare fresh clinopyroxene relics can be found and analyzed and, again, their compositions are characteristic for intra-oceanic fore-arc regions (Fig. 12.22b).

12.6.2 Alteration and Metamorphism

The mantle sequences of the ED ophiolites preserve evidence of a variety of post-magmatic fluid interaction processes that sample a range of temperatures, pressures, and fluid compositions. Alteration and metamorphism may have occurred at the ocean floor, below the oceanic crust, during and after tectonic emplacement, or upon recent exposure. Hence, a variety of origins and compositions have been proposed for metamorphic fluids affecting the ANS ophiolites (e.g., Azer 2013; Gahlan et al. 2018): (i) mantle-derived CO₂-bearing fluids (e.g., Boskabadi et al. 2017; Hamdy and Gamal El Dien 2017), (ii) seawater at near-bottom conditions (˂ 100 °C) (e.g., Snow and Dick 1995; Li and Lee 2006); (iii) both H₂O-rich and CO₂-rich fluids released from various layers of a subducting slab (e.g., Bostock et al. 2002; Hamdy et al. 2013); and (iv) hydrothermal fluids infiltrating during and after exhumation (˃100 °C) (e.g., Seyfried and Dibble 1980; Li and Lee 2006; Hamdy and Lebda 2007).

The most notable metamorphic process, affecting all the ultramafic rocks in the ED ophiolites, is serpentinization, a continuous process of replacement of anhydrous primary silicate minerals by hydrous phyllosilicates (e.g., O’Hanly 1996). Serpentinization reactions have been studied by (1) analysis of vent fluids from peridotite-hosted hydrothermal systems (e.g., Douville et al. 2002), (2) experiments and theoretical considerations (e.g., Allen and Seyfried 2003), and (3) petrographic studies of altered rocks (e.g., Evans 1977; Mével 2003).

The serpentinization of peridotites can occur in low-temperature (≤250 °C) and high-temperature (≥250 °C) environments (Evans 2010). At low temperature, olivine relics tend to retain their original composition, while high-temperature serpentinization allows interdiffusion of Mg and Fe diffusion, resulting in low Mg# olivine relics coexisting with serpentine. Furthermore, experimental studies have demonstrated that pyroxenes react faster than olivine at temperature above 250–300 °C, but olivine reacts faster than pyroxene at temperature <250 °C (e.g., Martin and Fyfe 1970; Janecky and Seyfried 1986; Allen and Seyfried 2003). In the ED case, olivine appears to have reacted most readily, followed by orthopyroxene and clinopyroxene, and the fresh olivine relics have high Mg# >88. Both features imply low-temperature hydration.

Serpentinization can be isochemical, which requires a significant volume increase and causes fracturing, or allochemical, which can in principle proceed at constant volume as fluids transport Mg²⁺, Fe²⁺, Ca²⁺, and Si²⁺ ions out of the system. Johannes (1969, 1970) provided benchmark experiments of equilibria in the system MgO–SiO₂–H₂O–CO₂ from 0.2 to 1.0 GPa pressure that demonstrates that serpentine coexists only with CO₂-poor fluid phases.

Upon prograde metamorphism, the reverse of serpentinization, i.e., dehydration, can occur, producing metamorphic olivine (e.g., Arai 1975; Evans 1977; Nozaka 2003). Such metamorphic olivine typically lacks chemical zonation or evidence of ductile deformation such as undulatory extinction (e.g., Mercier and Nicolas 1975). Metamorphic olivine has been reported in peridotites from ED ophiolites (e.g., Khalil and Azer 2007; Gahlan 2006) and attributed either to the thermal effects of granitoid intrusions (e.g., Gahlan and Arai 2009; Ahmed et al. 2012b) or to low-pressure high-temperature regional metamorphism (Gahlan et al. 2015a).

During or after serpentinization, many of the ED ophiolite ultramafic rocks were variably deformed and altered to mixtures of serpentine, talc, chlorite, carbonates, and magnetite (e.g., Ghoneim et al. 2003; Farahat 2008; Abdel-Karim et al. 2018; Azer et al. 2019; among others). The timing and fluid sources for events such as carbonatization, listwaenitization, and rodingization are controversial. Recently, Azer et al. (2019) distinguished two stages of carbonation in ED ophiolites. The first stage formed magnesite masses during deep-seated metasomatism and serpentinization, whereas the second stage emplaced carbonate veins after serpentinization, during obduction of the ophiolite. Likewise, listwaenitization in ED ophiolites also took place in two main metasomatic stages, the first associated with the formation of the oceanic crustal section and the second during obduction (Gahlan et al. 2018).

Rodingite is a massive light-colored rock composed of Ca-rich minerals produced by infiltration of Ca-bearing solutions into aureoles around serpentinized ultramafic bodies (Best 2003). Determination of the physical and chemical conditions of rodingite formation, concomitant with serpentinization, can provide detailed information on the effects of ancient seafloor hydrothermal processes and on the tectonic history of an ophiolite. Only a few studies have published data about ED rodingites (Takla et al. 1992; Abdel-Karim 2000; Surour 2019). Although the ED ophiolites have been intensively studied over the last two decades, the associated rodingites have not yet received much attention.

The metamorphic grade implied by the mineral assemblages in serpentinized ultramafic rocks of the ED ranges from greenschist facies (olivine–opx–serpentine–ferritchromite–magnetite) to amphibolite facies (olivine–opx–anthophyllite–talc–tremolite–antigorite–ferritchromite–magnetite) (e.g., Evans and Trommsdorff 1974; Evans 1977; Gahlan et al. 2015a; Obeid et al. 2016). The formation of ferritchromite rims around fresh chrome spinel cores, in particular, indicates greenschist or lower amphibolite facies prograde metamorphism (Evans and Frost 1975; Suita and Streider 1996; Mellini et al. 2005; Azer et al. 2019) with peak temperatures >500 °C and oxidizing conditions (e.g., Farahat 2008; González-Jiménez et al. 2009).

12.7 Proterozoic Versus Phanerozoic Mantle Sections: A Comparison

Paleoproterozoic (1650–2300 Ma) and Mesoproterozoic (1000–1400 Ma) ophiolites are less abundant globally than Neoproterozoic (543–1000 Ma) examples (Kusky 2004). The Dongwanzi ophiolite (2505 Ma) in north China is considered the world’s oldest, nearly complete and well-preserved Paleoproterozoic ophiolite complex (Kusky et al. 2001; Li et al. 2002; Kusky 2004; Huson et al. 2004). Other well-known Paleoproterozoic ophiolites include the Jormua ophiolite (~2000 Ma) of Finland (Peltonen and Kontinen 2004), the Purtuniq ophiolite (1998 Ma) in the Cape Smith Belt of Canada (Scott et al. 1992), and the unique Payson ophiolite (1730 Ma) in the USA (Dann 2004). Recently, some Mesoproterozoic ophiolites have been described from the Karelian Shield of West Africa and from the southwestern USA (e.g., St-Onge et al. 1989; Abouchami et al. 1990; Boher et al. 1992; Dann 2004). By contrast, the youngest ophiolite complex pertaining to the Neoproterozoic era is the Agardagh Tes–Chem ophiolite (570 Ma), Tuva, Central Asia (Pfänder and Kröner 2004). The Arabian–Nubian Shield and the ED of Egypt in particular are rich in well-preserved and exposed Neoproterozoic ophiolites (e.g., Stern et al. 2004 and references therein).

Highly depleted harzburgites with high-Cr# (≥0.7) spinel are very common in Proterozoic ophiolites (e.g., Quick 1990; Liipo et al. 1995; Vuollo et al. 1995; Ahmed et al. 2001; Gahlan et al. 2015b; among others). However, similar rocks are not at all uncommon in Phanerozoic ophiolites (e.g., England and Davies 1973; Arai 1997; Tamura et al. 1999; Ishiwatari et al. 2003 and references therein). Likewise, spinel with Cr# as high as 0.8 is commonly found in modern fore-arc regions (e.g., Marian Trench; Ohara and Ishii 1998).

In terms of chromian spinel chemistry, the Neoproterozoic ED ophiolitic peridotites are similar to a number of published Neoproterozoic examples across the Red Sea in the Arabian Shield (Stern et al. 2004 and references therein) (Fig. 12.20b). Like the Arabian Shield peridotites, the ED peridotites are best interpreted in an SSZ setting (e.g., Arai and Yurimoto 1995; Stern et al. 2004).

Overall, the association harzburgite–dunite–chromitite in the Neoproterozoic ED ophiolites and in Phanerozoic cases is similar in their field, petrographic and mineral chemical characteristics (Fig. 12.22). The formation process of podiform chromitites in ED ophiolites is therefore probably similar to that in Phanerozoic cases, namely, by melt–peridotite interaction and subsequent melt mixing in an SSZ setting (e.g., Ahmed et al. 2001). Comparing the spinel chemistry of the Neoproterozoic ED chromitites with a compilation of Phanerozoic equivalents, we find that the most notable difference is modestly but systematically lower TiO₂ in the ED cases. This can be attributed to somewhat higher degrees of partial melting or a more fractional character of melting in development of the Neoproterozoic examples (Kelemen et al. 1995; Asimow and Stolper 1999). In general, however, the Neoproterozoic ophiolites support the notion that subduction processes have on the whole been unchanged over the last billion years of Earth history. No dramatic differences stand out that require a non-uniformitarian interpretation of these rocks.

12.8 Mineralization

The mantle sections of ophiolites are worthy targets for mining and exploration. They host a variety of ores (chromite, gold, and iron–nickel laterites) and industrial minerals (talc, asbestos, and serpentine) (e.g., Coleman 1977; Klemm and Klemm 2013; Gahlan et al. 2018; Fu et al. 2019; among others). In the ED, in particular, there is a strong association between ophiolites and mineralization, including chromite, talc, asbestos, platinum-group elements, Cu–Ni–Co, magnesite, and gold (Klemm et al. 2001; Kusky and Ramadan 2002; Ahmed and Hariri 2008; Azer et al. 2019). We will discuss some of the major resource types in the following sections.

12.8.1 Chromitite

Chromite deposits in Egypt are commonly hosted by serpentinized ultramafic rocks, widely distributed in the central and southern ED. In most cases, chromitites occur as lenticular bodies of variable size, from thin pencil-shaped layers (a few centimeters long) to large pods (up to 30 m along strike). The average pod is relatively small (≤5 m length and ≤2 m width). The podiform chromitites (Thayer 1964) are not especially abundant in the ED mantle sections as a whole and are commonly concentrated in the shallowest parts, closest to the petrologic Moho (e.g., Hume 1937; Amin 1948; Takla et al. 1975; Takla and Noweir 1980; El Haddad and Khudeir 1989; Ahmed et al. 2001; Gahlan et al. 2015b; among others). They are commonly concordant to sub-concordant with the plano-linear fabrics of the host rocks and enveloped by dunite those grades outward to harzburgite (Fig. 12.2). Ore pods range massive, nodular, anti-nodular, and disseminated textures are observed.

The ED chromitites are mostly of metallurgical-grade (Cr₂O₃ > 40 wt% and Al₂O₃ < 20 wt%). Notably, the South ED chromitites are more refractory (i.e., Cr-rich) than the Central ED examples. Arai (1997) concluded that size of chromitite pods is highly dependent on the chemistry of the host peridotites, particularly spinel Cr#. Globally, the largest chromitite pods are hosted by moderately refractory peridotite host (spinel Cr# = 0.4–0.6), whereas only very small chromitite pods, if any, are typically hosted either by cpx-free highly refractory peridotites (spinel Cr# > 0.7) or by lherzolitic peridotites (spinel Cr# ≤ 0.3). As the ED host peridotites are all of the highly refractory variety, the low overall size and abundance of podiform chromitites in the ED mantle sections appear consistent with the tendency proposed by Arai (1997). Like podiform chromitites worldwide, those of the ED mantle section are well explained as the result of melt–peridotite interaction and subsequent melt mixing in the SSZ upper mantle (e.g., Lago et al. 1982; Paktunc 1990; Arai and Yurimoto 1994, 1995; Zhou et al. 1994, 1996; Arai 1997; Ahmed et al. 2001; Ahmed 2013; Azer et al. 2019).

12.8.2 Gold

Prospects and productive gold mining sites are widespread in the ED mantle sections (Fig. 12.23), particularly in the carbonatized serpentinites (e.g., Botros 2002, 2004; Zoheir and Lehmann 2011; Abd El-Rahman et al. 2012a; Boskabadi et al. 2017; Gahlan et al. 2018). Gold mining in these areas extends back to Pharaonic times (Harraz 2000; Klemm and Klemm 2013). According to Botros (2004), the El-Sid gold deposit is a good example of vein-type mineralization hosted in sheared ophiolitic ultramafic rocks. The El-Sid gold mine is confined to hydrothermal quartz veins at the contact between granite and serpentinite except for extensions into the serpentinite along a thick zone of graphite schist (El-Bouseily et al. 1985). Historically worked gold deposits are associated with ophiolitic rocks in the Wadi Allaqi region.

Fig. 12.23

Comparison between Proterozoic and Phanerozoic ophiolitic peridotite (P), dunite (D), and chromitite (Ch), in terms of spinel composition. The Phanerozoic ophiolite data are from Arai (1997) and references therein; the Kenticha Hill chromitite data are from Bonavia et al. (1993), the Outokumpu Jormua ophiolite data are after Liipo et al. (1995), and Egyptian ED data are from Ahmed et al. (2001) and Gahlan (2015a)

During serpentinization and metamorphism, gold behaves, to some extent, like the fluid-mobile elements (B, Li, S, As, Rb, Sr, Sb, Cs, Ba, Pb, and U) (e.g., Deschamps et al. 2013; Gahlan et al. 2018; and references therein). Typically, gold concentrations are ~1 ppb in depleted mantle (Salters and Stracke 2004), about 2.8 ppb in average ophiolitic rocks (Crocket 1991), and ~3–5 ppb in serpentinites (Buisson and Leblanc 1987). The Fawakhir goldmine shows gold concentrations up to ~30 g/ton (Hussein 1990), supporting the suggestion of Buisson and Leblanc (1987) that Neoproterozoic ophiolitic mantle sections are modestly enriched in gold. Abd El-Rahman et al. (2012b) noted the dominance of vein-type gold deposits along the eastern side of the NW-trending Najd Fault System and attributed this pattern to the abundance of serpentinite bodies in the arc–fore-arc belt.

A spatial and genetic relationship has been observed between carbonated ultramafic rocks, subsequent granite intrusions, and gold mineralization. Listvenites are considered a prime target for gold prospecting in ophiolites (Botros 2004; Ahmed and Hariri 2008). For example, the listvenites at G. Sirsir show gold concentrations up to 6584 ppb (as well as As and Ag enrichment) (Gahlan et al. 2018). Azer (2008) extended this association into a genetic link between gold mineralization and carbonate alteration of ED ophiolites, whereby carbonatization preconcentrates gold up to 1,000 times compared to the original ultramafic rocks and interaction with hydrothermal systems associated with granite intrusions further concentrates and remobilizes gold (Cox and Singer 1986; Azer 2008). However, the relation between Au en-richment and the carbonatization process is controversial (e.g., Azer 2013; Boskabadi et al. 2017; Gahlan et al. 2018; and references therein).

A variety of potential minerals has been proposed to host gold in serpentinites including nickel sulfides (e.g., Takla and Surour 1996; Khalil et al. 2003; Emam and Zoheir 2013), Cu–Fe–Ni sulfides (e.g., Ferraris and Lorand 2015), olivine (Au as nano-inclusions), and lizardite (Boskabadi et al. 2017). Breakdown of these Au-bearing minerals upon infiltration of CO₂-bearing hydrothermal fluids liberates fluid-mobile elements and results in an Au-bearing CO₂-rich fluid phase. The source of CO₂-rich hydrothermal fluid phase has been a matter of debate, but a mixture of mantle-derived and surface-derived CO₂ has been widely accepted (Boskabadi et al. 2017; Gahlan et al. 2018; and references therein).

It has been proposed that the contrast in rheology and permeability between listvenites and its country rocks then promotes precipitation of gold at their interfaces (e.g., Zoheir and Lehmann 2011; Azer 2013; Zoheir and Moritz 2014; Boskabadi et al. 2017; Gahlan et al. 2018). This mechanism is most favored at greenschist facies metamorphic conditions, where brittle–ductile deformation is available to form pathways for gold mineralization (El-Gaby et al. 1988; Botros 2004). The result is development of auriferous quartz veins within sheared serpentinites and listvenites (e.g., Zoheir and Lehmann 2011; Abd El-Rahman et al. 2012b). The auriferous veins are composed mainly of quartz and carbonates with subordinate sulfides (mainly pyrite and chalcopyrite) (Botros 2002, 2004). Gold occurs either as native, tiny specks within quartz veins or as inclusions in sulfides, particularly in pyrite and arsenopyrite (Abd El-Rahman et al. 2012a).

12.8.3 Talc–Carbonates

Worldwide production of talc has increased from 7.4 million tons in 2011 to about 10 million tons in 2018. The major producing countries are China, USA, Finland, and France. Egypt produced 12,924 tons of talc in 2011 and about 172,181 tons in 2015. The Darhib and Atshan deposits have been the most intensively exploited to date. Talc was mined and used by ancient Egyptians for beads and for cosmetic vessels in both the Middle and New Kingdoms. High-purity talc (>95%) is used in cosmetics, steatite, cordierite ceramics, paper, and plastics. Medium purity talc (75–95%) is used in paper, plastics, wall tiling, paint, and rubber. Low-purity talc (e.g., <75%) is used in paint, roofing materials, flooring, and fertilizers.

Talc–carbonate rocks are widely distributed at various ED localities. Economically significant deposits of talc are already being mined along the suture in the Wadi Allaqi region, but huge quantities of practically untapped talc–carbonate rocks could serve as an important potential source of magnesia for strategic industries such as refractory ceramics and electrical insulators. There are two types of talc–carbonate deposits in the ED: (1) talc–carbonates associated with volcano-sedimentary rocks of island-arc stage, and (2) talc–carbonates associated with ophiolitic ultramafic rocks. Talc–carbonates associated with ophiolitic rocks are well developed in the Barramiya, Atud, and Atalla areas. These rocks were called Barramiya rocks (Hume 1934; Rittmann 1958). They occur either as separate bodies or associated with serpentinite. They are developed along major faults and shear zones cutting serpentinite units or at thrust contacts between obducted serpentinites and other country rocks of island-arc affinity.

As discussed above, talc–carbonate rocks associated with the serpentinites are easily recognized by their distinct appearance in the field. A systematic campaign is needed to map the occurrences of talc–carbonates in Egypt and to classify them according to standard industrial specifications in order to locate the highest-purity deposits (suitable for ceramic applications) and to define Egypt’s reserves of talc.

12.8.4 Magnesite

Reaction between CO₂-rich fluids and olivine or serpentine usually produces the industrially useful mineral magnesite (MgCO₃) (e.g., Klein et Garrido 2011), which is commonly found worldwide in networks of veins in the ultramafic parts of ophiolite sequences. In the ED, both massive pods of magnesite and very high-purity magnesite veins have been described at numerous ophiolite-associated localities (Azer et al. 2019). The massive magnesites, as discussed above, are formed by replacement during early serpentinization processes, whereas the vein-type magnesite is formed later. The magnesite veins fill fractures and cavities in serpentinites (e.g., Salem et al. 1997; Ghoneim 2003; Ghoneim et al. 1999).

12.8.5 Serpentinite

Serpentinite itself is a decorative and industrial rock that can be used, for example, for neutron shielding in nuclear reactors (e.g., Abulfaraj and Kamal 1994). The mineralogy, color, and mechanical properties determine whether given serpentinites have decorative or industrial value. Streaks of associated minerals such as talc and carbonates for attractive visual patterns that are considered controlling factors in the value of serpentinites as ornamental stones (e.g., Ismael and Hassan 2008). They are many serpentinite quarries in the ED; the largest one is the Wadi Sodmein quarry (Figs. 12.24 and 12.25).

Fig. 12.24

Simplified geological map showing the extent of the ANS (inset), the distribution of ophiolitic bodies, and selected Au deposits in the Egyptian segment of ANS

Fig. 12.25

Serpentinite quarry in Wadi Sodmein showing cut surfaces of Atg-serpentinite

12.9 Conclusions

• The mantle sections or lower units of ophiolites in the Eastern Desert of Egypt generally form massive ridges and sheet-like bodies of serpentinized harzburgite, dunite, and rarely lherzolite with chromitite pods and ultramafic cumulates. They are fault-bounded and often well preserved in the troughs of major synforms. The contacts between the mantle and crustal sections of these ophiolites were originally magmatic but are now inevitably disrupted by tectonism.

• Petrographic examination shows that almost all primary silicates are typically replaced, except for chromian spinel and rare relics of olivine and sometimes pyroxene.

• Regional metamorphism to greenschist or amphibolite facies and CO₂-metasomatism resulted in an array of talc–carbonate, listvenite, magnesite, and carbonate-bearing meta-ultramafic rocks.

• The modal mineralogy (i.e., low or zero abundance of clinopyroxene), mineral chemistry (high-Cr#, low-TiO₂ spinel; high Fo and NiO in relict olivine; low Al2O3 in relict orthopyroxene), and whole-rock chemistry (high-Mg#; enrichment of Ni, Cr, and Co; depletion of Al₂O₃, TiO₂, and CaO; depleted and unfractionated REE patterns) of the ED mantle rocks all point to highly refractory protoliths that experienced high degrees of partial melt extraction. In all these regards, they resemble modern fore-arc peridotites.

• The ED podiform chromitites resemble their Phanerozoic equivalents and may have been formed by the same mechanisms. In general, the ED Neoproterozoic ophiolite mantle sections closely resemble Phanerozoic fore-arc equivalents, suggesting little change in the geothermal and tectonic regime of Earth’s subduction zones since the Late Proterozoic era.

• A variety of economically viable deposits of ore (e.g., chromite and gold) and industrial (e.g., asbestos, talc, magnesite, and serpentinite) minerals are spatially associated with ED mantle sections.