The Egyptian Nubian Shield Within the Frame of the Arabian–Nubian Shield

Abd El-Wahed, Mohamed; Hamimi, Zakaria

doi:10.1007/978-3-030-49771-2_2

Mohamed Abd El-Wahed⁸ &
Zakaria Hamimi⁹

Part of the book series: Regional Geology Reviews ((RGR))

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7 Citations

Abstract

Gondwana Supercontinent in eastern and southern Africa formed by collision and amalgamation of two crustal plates, provisionally named East Gondwana and West Gondwana and the Mozambique Ocean intake between 841 and 632 Ma. East Gondwana consists of the Arabian–Nubian Shield (ANS) and the older crystalline basement in Madagascar, India, Antarctica and Australia, while West Gondwana consists of much of Africa and South America. Collision and amalgamation of East and West Gondwana formed the East African Orogeny. The supercontinent of Gondwana ranged from Neoproterozoic (~550 Ma ago) to Carboniferous (~320 Ma ago). Gondwana became the largest continental crust during the Paleozoic Era (~100 million km²). During the Carboniferous, Gondwana amalgamated with Euramerica resulting in the formation of the supercontinent, Pangaea. Three orogeneses were recognized during the 1990s: the East African Orogeny (650–800 Ma), Kuunga Orogeny (including the Malagasy Orogeny in southern Madagascar) (550 Ma)—the collision between East Gondwana and East Africa in two steps—and the Brasiliano Orogeny (660–530 Ma)—the collision between South American and African Cratons. Formation of arcs in the ANS occurred over a ~300-million-year period including supercontinent Rodinia break-up and the assembly of supercontinent Gondwana. The ANS represents one of the best documented examples of Late Proterozoic to Early Paleozoic (950–450 Ma) crustal growth through processes of lateral arc–arc terrane accretion. The tectonic development of the ANS spans three phases spanning over 600 Ma: accumulation of arc terrains inside the Hijaz Magmatic Arc, accompanied by accretion of the Hijaz Magmatic Arc against the Nile Craton and reworking of the accreted arc after accretion. The Egyptian Nubian Shield (ENS) covers ~100,000 km², crops out along the Red Sea Hills in the Eastern Desert and southern Sinai, as well as limited areas in the south Western Desert (Oweinat area, 2673 ± 21 Ma). The ENE covers the northeastern part of the East African Orogeny and stretches over approximately 800 km parallel with the Red Sea coast between latitudes 22° 00΄ 00˝ and 28° 40΄ 00˝ N. The rocks are covered by Nubia sandstone, Miocene and later sediments in their western and eastern margins. The Eastern Desert of Egypt is divided into three domains, namely, the northern Eastern Desert (NED), central Eastern Desert (CED) and southern Eastern Desert (SED); these domains were developed in different tectonic settings and show a characteristic younging from (SED) to (NED). Geologically, gneisses, migmatites and schists dominate the SED as the oldest units, and are followed by ophiolites, volcanic arc lithologies and granitoid plutons. The amount of ophiolites increases and forms with the arc metavolcanics the main types in the CED. The ophiolites and the metavolcanics are occasionally unconformably and tectonically overlained by Dokhan volcanics and molasse sediments. The older gneisses and migmatites form prominent domal structures (e.g. Meatiq, Sibai, Hafafit, El-Shalul). Syn-tectonic and late tectonic granitoids are also present. The NED is characterized by younger rocks, such as Gattarian granites, Dokhan volcanics and Hammamat molasse sediments, whereas older rock types rarely occur. The bulk of the crust of the SED was created prior to 650 Ma, while the major pulses of the CED occurred in the interval (685–575 Ma). In the NED and Sinai, the crust was principally formed in the period (625–575 Ma).

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The Arabian–Nubian Shield, an Introduction: Historic Overview, Concepts, Interpretations, and Future Issues

The Southern Part of the Arabian–Nubian Shield in Kenya and Tanzania

The Arabian–Nubian Shield Within the Neoproterozoic Plate Tectonic Circuit

2.1 Formation and Amalgamation of Gondwana Supercontinent

Gondwana is the name for the southern half of an ancient supercontinent known as Pangaea or Pangea that existed some 300 Ma, along with a northern supercontinent known as Laurasia (Meert et al. 2011). The continent broke up about 180 Ma and eventually split into landmasses we recognize today: South America, Africa, Madagascar, Sri Lanka, India, the Arabian Peninsula, Antarctica and Australia (Fig. 2.1). Gondwana was the largest unit of continental crust on Earth for more than two hundred million years. The name “Gondwana” or “Gondwanaland” is derived from a tribe in India (Gonds) and “wana” meaning “land of”. This name was first used in 1879 by Medlicott and Blanford from the Indian Geological Survey. Wegener (1915) enlarged the concept of Gondwana and postulated his ideas regarding continental drift, and the existence of a previously united Gondwana. Wegener once believed that all of the continents were together in a “Urkontinent” before splitting apart and moving to their present places. Most of Wegener’s findings on fossils and rocks are right, he was in certain ways highly incorrect. For instance, Wegener figured that the continents should have plowed like icebreakers into the ocean crust. Plate tectonics is now the widely accepted theory that Earth’s crust is fractured into rigid, moving plates.

Several events, collectively known as the Pan-African orogeny, led to the amalgamation of most of the continental fragments of a much older supercontinent, Rodinia. Rodinia is a supercontinent that was united 1.3–0.9 billion years ago during Grenville Orogeny and broke up 750–633 million years ago (McMenamin and McMenamin 1990; Li et al. 2008; Meert 2012). It was surrounded by an ocean called Mirovia. Valentine and Moores (1970) were perhaps the first to identify a Precambrian supercontinent called Pangaea that was renamed as “Rodinia” by McMenamin and McMenamin (1990). The fragments of the break-up of Columbia collided and were assembled by global-scale 2.0–1.8 Ga collisional events and formed Rodinia at c. 1.23 Ga (Zhao et al. 2002, 2004). Condie (2002) reported that Rodinia was created between 1300 and 900 Ma and that the continental disintegration happened between 950 and 600 Ma (Fig. 2.2). The break-up of Rodinia started around 950 Ma ago and persisted until ca. 600 Ma (Condie 2002; Rogers and Santosh 2004).

The Mozambique Belt is one of the 800–650 Ma orogenic belts and was initially described as the suture dividing East from West Gondwana. The Mozambique Ocean separated the Congo–Tanzania–Bangweul Block of central Africa from Neoproterozoic India (India, the Antongil Block in far eastern Madagascar, the Seychelles, and the Napier and Rayner Complexes in East Antarctica). The Azania continent was an island in the Mozambique Ocean and include much of central Madagascar, the Horn of Africa and parts of Yemen and Arabia (Collins and Pisarevsky 2005). The final formation of Gondwana occurred about 500 million years ago (Figs. 2.3, 2.4 and 2.5). Gondwana is superficially divided into western half (South America and Africa) and an eastern half (Madagascar, Antarctica, Sri Lanka, India and Australia) and Ediacaran–Cambrian age by coalescence of East Gondwana with West Gondwana with closing of the Mozambique ocean and development of the Mozambique Belt (Stern 1994a, b; Meert et al. 2011; Johansson 2014).

Gondwana amalgamated with Euramerica during the Carboniferous to create a broader supercontinent called Pangaea. During the Mesozoic era, Gondwana (and Pangaea) were gradually fragmented. During Silurian, Gondwana supercontinent (Australia, Antarctica, India, Arabia, Africa, and South America, Florida, southern Europe, and the Cimmerian terranes, namely, Turkey, Iran, Afghanistan, Tibet and the Malay Peninsula) was centred over the South Pole. Rheic Ocean is an east–west ocean separating the southern European sector of Gondwana from northern Europe (Baltica) and was basically a southwestern extension of the Paleotethys Sea. During the Caledonian orogeny (410 Ma ago), a minor supercontinent named Euramerica was created due to collision between the Avalonia cratons, Baltica and Laurentian. The Caledonian orogeny was a mountain-building era recorded in the northern parts of Ireland and Britain, the Scandinavian Mountains, Svalbard, eastern Greenland and parts of north-central Europe (Torsvik and Cocks 2013). In the Permian, the supercontinent of Euramerica became part of the major Pangaea supercontinent. Euramerica became part of Laurasia in the Jurassic when Pangaea split into two separate continents, Gondwana and Laurasia. Some 180 Ma, in the Jurassic Period, the western half of Gondwana (Africa and South America) separated from the eastern half (Madagascar, India, Australia and Antarctica) (Figs. 2.6 and 2.7). In the Cretaceous, Africa broke away from South America, leading to opening of the South Atlantic Ocean at about 140 Ma. In the same period, the central part of the Indian Ocean opened, and India parted from Madagascar, and Australia slowly rifted away from Antarctica. Also, Laurasia divided into the North America and Eurasia continents. Baltica was part of Eurasia, and the Laurentian craton was part of North America while Avalonia became split between the two. The remains of Gondwana constitute two-thirds of the continental region of today (Torsvik and Cocks 2013). India collided with Eurasia creating the Himalayan Mountains 50 million years ago, when the northward-moving Australian plate has only started its collision on the southern edge of south-east Asia, a collision still ongoing today.

There are basically two geodynamic scenarios to explain the formation of the Gondwana Supercontinent in eastern and southern Africa during the Pan-African orogenic cycle (Westerhof et al. 2008). The first scenario involved collision and amalgamation of East Gondwana and West Gondwana (~640 to ~530 Ma) (Shackleton 1994; Kröner et al. 2001; Jacobs et al. 2006; Westerhof et al. 2008) and the Mozambique Ocean consumed between 841 and 632 Ma (Cutten and Johnson 2006). In this model, East Gondwana comprised juvenile crust (ANS, and older crystalline basement in Australia, Antarctica, India and Madagascar), and West Gondwana was composed of most of South America and Africa (Figs. 2.4, 2.5 and 2.7). Collision and amalgamation of East and West Gondwana produced the East Africa Orogen (EAO) (Stern 1994a, b) that it is an N–S directed fold belt (6000 km). This orogenic cycle dated ~650–490 Ma (Cahen and Snelling 1966). Kröner (2006) reported peak granulite facies metamorphism in East Africa and Madagascar between 640 and 550 Ma. The second scenario assumes collision and amalgamation of East Gondwana (ANS and older crystalline basement of the Dharwar Craton of southern India, Madagascar and the eastern granulites of Kenya and Tanzania), West Gondwana (Central Africa craton) and South Gondwana (Antarctica and the Kalahari craton) (Grantham et al. 2003) (Fig. 2.1). South Gondwana has been integral since the Grenvillian Orogeny at ~1.0 Ga. In this scenario, the EAO suture directed by the N–S implies Kuunga Orogen suture directed to the E–W, comprising from W to E, the Damara–Lufilian–Zambezi (DLZ) Belt (Fig. 2.8), the Lúrio Thrust Belt (LTB) and, further eastwards, thrust belts of Sri Lanka (Westerhof et al. 2008). In this scenario, pan-African remobilization of Mesoproterozoic and older crust south of the Kuunga Suture is confined to metamorphic overprinting throughout linear N–S-directed shear zones (Manica Shear Zone, Fig. 2.1) along the eastern margin of the Zimbabwe Craton (Manhiça et al. 2001; GTK Consortium 2006; Westerhof et al. 2008).

Collision and amalgamation of the Central Africa Craton (West Gondwana) and Kalahari Craton (South Gondwana) and Closure of the Zambezi–Adamastor Oceanic Basin (Johnson et al. 2005) resulted in the Damara–Lufilian–Zambezi Belt (Fig. 2.8), a major Neoproterozoic suture (Burke et al. 1977; Oliver et al. 1998; John et al. 2003; Johnson and Oliver 2000, 2004). The Lufilian and Zambezi segments of the Damara–Lufilian–Zambezi Belt can be described as a thin- and thick-skinned, double-verging orogen with thrust transport to the SSE (Müller et al. 2001; Kuribara et al. 2018) and NNE–NE (Hanson et al. 1994; Johnson and Oliver 2004), suggesting oblique convergence between West and South Gondwana (Westerhof et al. 2008).

2.2 Orogens that Shaped the Greater Gondwana

The assembly of continental blocks forming West Gondwana began within the 900–700 Ma time frame, and reached completion at 550–530 Ma, based on paleomagnetic, geologic and isotopic data (Meert and Van Der Voo 1997; Rogers and Santosh 2004; Teixeira et al. 2007). Three orogenies were recognized during the 1990s: (i) The East African Orogeny (650–800 Ma), (ii) Kuunga Orogeny (including the Malagasy Orogeny in southern Madagascar) (550 Ma) and (iii) The Brasiliano Orogeny (660–530 Ma). The East African and Kuunga are the resultant of the collision between East Gondwana and East Africa, whereas the Brasiliano Orogeny is the result of the collision between South American and African Cratons (Meert and Van Der Voo 1997).

2.2.1 The West- African–Brasiliano Orogeny

Kennedy (1964) identified a major period of deformation and metamorphism in Africa that occurred at approximately 500 Ma. Orogeny at this time was soon recognized to be widespread throughout Africa and was referred to as ‘‘Pan-African”. Geologists in South America also found numerous orogenic belts of the same age and named them Brasiliano. The descendants of Rodinia have evolved during the whole Neoproterozoic and the Cambrian, and they interacted during that time (ca. 410 Ma as a whole) and generated the so-called Brasiliano/Pan-African Mobile Belts (Figs. 2.9 and 2.10) (Brito-Neves et al. 1999). Brasiliano Orogeny refers to a sequence of Neoproterozoic orogenies found mostly in Brazil, and elsewhere in South America. The Brasilia Belt is a complex orogenic system (thin and thick skin), 1100 km long, displaying structural vergence towards the São Francisco Craton, with P and T metamorphic conditions increasing progressively towards the west (Brito-Neves et al. 1999). The Brasiliano Orogeny is a common term for Pan-African/Brasilian Orogeny that extended not only in South America but across most of Gondwana (Rogers and Santosh 2004; Kröner and Stern 2004). The orogeny resulted in the closing of several oceans including the Puncoviscana, the Peri-Franciscano, the Goianides and the Adamastor (Tohver et al. 2010; Frimmel and Hartwig 2010). The West African–Brasilian Orogeny had been tectonically active until the end of the Cambrian period (~490 Ma; Schmitt et al. 2004; Meert and Lieberman 2008). Prior to Gondwana break-up, the Congo Craton originally linked with the Sao Francisco Craton forming a larger continental mass wrapped by Brasiliano orogenic belts (Trompette 1994, 2000). In southern Africa, the Congo and Kalahari Cratons are separated by a transcontinental Pan-African orogeny comprising the Damara Belt, the Lufilian Arc and the Zambezi Belt (Figs. 2.8 and 2.9) (Kuribara et al. 2018). Brasiliano Orogeny along the western margin of southern Africa comprises the West Congo, Kaoko, Gariep and Saldania Belts (Figs. 2.9 and 2.10) (Hanson 2003).

2.2.2 The Damara–Zambesi–Lufilian Orogeny

The Damara–Zambezi–Lufilian Orogeny developed essentially by closure of linked, narrow ocean basins (Hanson 2003). The Zambezi Belt (Figs. 2.8 and 2.9) lies between the Congo Craton and the Kalahari Craton and branches off the Mozambique Belt in northern Zimbabwe and southern Zambia (Johnson and Oliver 2002; Kröner and Stern 2004; Rogers and Santosh 2004; Kuribara et al. 2018). The Kalahari Craton is an old and stable part of the continental lithosphere that consists of the Kaapvaal, the Zimbabwe Craton, the Limpopo Belt and the Namaqua Belt, and occupies a large portion of South Africa (Johnson and Oliver 2002). The Congo Craton is an ancient Precambrian craton covered by the Palaeozoic to recent. Congo Basin together with Tanzania, West African, Zimbabwe and Kaapvaal cratons constitutes the recent continent of Africa (Johnson and Oliver 2002). These four cratons were formed between about 3.6 and 2.0 Ga and have been tectonically stable since that time (Johnson and Oliver 2002).

The Zambezi Belt was developed through two tectonothermal events, one between about 890 and 880 Ma and the other about 550 and 520 Ma. Both events affected the existing Archean to Mesoproterozoic rocks (Johnson and Oliver 2002; Kuribara et al. 2018). The second event was produced during the assembly of the Gondwana Supercontinent as a result of the Congo–Kalahari collision at the end of the Neoproterozoic (Johnson and Oliver 2002; Hargrove et al. 2003).

The Damara Orogeny occurred late in the creation of Gondwana, and the Damara Belt is exposed mainly in Namibia between the Kalahari and the Congo cratons (Jung and Mezger 2003; Gray et al. 2008). Damara Belt (Figs. 2.8, 2.9 and 2.10) is produced by the closing of the Damara and Adamastor oceans and continues northwards into the Kaoko Belt and southwards into the Saldania and Gariep Belts (Kröner and Stern 2004). The Damara Orogeny involved the suturing of the Congo–São Francisco and Río de la Plata Cratons at 580–550 Ma (together with India forming northern Gondwana) before the amalgamation of the Kalahari and Mawson Cratons in the Kuunga–Damara Orogeny at 530 Ma (southern Gondwana) (Jung and Mezger 2003; Gray et al. 2008). The Adamastor Ocean closed southwards from the Araçuaı́ Belt (São Francisco Craton, now in South America) to the Kaoko Belt (Congo Craton, now in Africa) 580–550 Ma and 545–530 Ma Kalahari Craton in southern Africa (Gray et al. 2008). The Kaoko Belt is developed by closing the Adamastor Ocean and comprises a shear zone in southern Angola known as the 733–550 Ma-old Puros lineament (Kröner and Stern 2004). It branches north-west from the Damara Belt into Angola and contains strongly deformed rocks (2030–1450 Ma-old). No island arcs or ophiolite are known from the Kaoko Belt (Kröner and Stern 2004). The peak of deformation and metamorphism in Damara Orogeny occurred at 530–500 Ma, while thrusting against the Kalahari Craton occurred at 480 Ma (Jung and Mezger 2003; Gray et al. 2008). All African cratons were assembled by c. 550 Ma, and the final amalgamation of north and South Gondwana were intra-cratonic during the last stages of the Damara–Kuunga Orogeny (Gray et al. 2008). The Damara Orogeny produced the Naukluft Mountains in central Namibia from 550 Ma to 495 Ma (Gray et al. 2008).

The Lufilian Arc (Fig. 2.8) is about 800 km long that extends across eastern Angola, the Katanga Province of the southern Congo and the north-west of Zambia (Ray et al. 2010; Laznicka 2010; Ray et al. 2010). It represents the continuation of the Damara Belt in Namibia northern Botswana (Kröner and Stern 2004). Its global economic importance comes from its enrichment in copper and cobalt deposits.

2.2.3 The East African Orogeny

The EAO (~6000 km N–S) represents the largest continuous Neoproterozoic–Cambrian orogeny on Earth (Fritz et al. 2013). It is subdivided into the ANS in the north, composed largely of juvenile Neoproterozoic crust (Stern 1994a, b, 2002; Johnson et al. 2011; Fritz et al. 2013) and the Mozambique Belt (MB) in the south (Figs. 2.1 and 2.11) comprising mostly pre-Neoproterozoic crust with a Neoproterozoic–early Cambrian tectonothermal overprint (Fritz et al. 2005; Collins 2006; De Waele et al. 2006; Bingen et al. 2009; Fritz et al. 2013).

In the early 1980s, the concept of Gondwana assembly in the Neoproterozoic via East and West Gondwana (Australia–India–Antarctica and Africa–South America) at the Pan-African Mozambique Belt (MB) was first proposed (McWilliams 1981; Stern 1994a, b). The main stage in the Late Precambrian Gondwana assembly and the continental collision between East and West Gondwana along the line of the Mozambique Belt was named the EAO (Stern 1994a, b). The Mozambique Belt is a suture in the earth’s crust between East and West Gondwana that extends from East Antarctica through East Africa (Tanzania, Kenya, Uganda, Rwanda, Burundi and South Sudan) up to the Arabian–Nubian Shield (Cutten 2002). Some of the components in the Mozambique Belt produced when the Mozambique Ocean opened and the other formed when the ocean later closed.

Bogdanova et al. (2009) concluded that Rodinia broke up in four stages between 825 and 550 Ma:

(i)
The break-up was superplume initiated (Large low-shear-velocity provinces, LLSVPs) around 825–800 Ma whose influences—including crustal arching, copious bimodal magmatism and thick rift-type sedimentary sequences—have been recorded in the Arabian–Nubian Craton, Tarim, India, South China, South Australia and Kalahari.
(ii)
Rifting proceeded in these same cratons (800–750 Ma) extending to Laurentia and maybe Siberia. India, Madagascar and the Congo–Säo Francisco Craton were split from Rodinia at this time or were never actual components of the supercontinent.
(iii)
As the central areas of Rodinia reached the Equator around 750–700 Ma, a new pulse of magmatism and rifting continued the disassembly in western Kalahari, West Australia, South China, Tarim and most margins of Laurentia.
(iv)
650–550 Ma several events coincided: the opening of the Lapetus Ocean; the closing of the Braziliano, Adamastor, and Mozambique Oceans; and the Pan-African Orogeny. The result was the formation of Gondwana.

The late Mesoproterozoic–early Neoproterozoic events of the southern segment of the East African Orogeny include separation of the Congo and the contiguous East Sahara Cratons from Rodinia (~1200 Ma), evolution of the sedimentary basin in Kenya (~820 Ma), development of a passive margin east of the Tanganyika shield of the Congo craton, and migmatization, ophiolite emplacement and metamorphism in the Kenyan segment of the East African orogen (Shackleton 1986; Mosley 1993; Unrug 1997). In the northern EAO, bimodal rift-related magmatism has been dated at 870–840 Ma in the Nubian shield and at 880 Ma in the Arabian shield (Stern 1994a, b; Unrug 1997). The rifting events in the late Mesoproterozoic–early Neoproterozoic include opening of large Arabian–Nubian and Pharusian Oceans east and west of the Congo–East Sahara–Nile craton cluster. At about 1000 Ma, passive margins developed along the eastern borders of the Rio de la Plata, Amazonian and West African cratons (Trompette 1994). The Adamastor Ocean opened between the Congo, Saõ Francisco, Rio de la Plata and Kalahari cratons (Powell 1993).

Two partly incomparable scenarios have been proposed for Gondwana assembly (Meert 2003). In the first model, the EAO developed from an accretionary orogeny involving the amalgamation of arcs and evolved into a collisional orogeny when the Neoproterozoic Azania Continent collided with the Congo–Tanzania–Bangweulu Block at c. 640 Ma (Collins and Windley 2002). In the other model, the East Gondwana assembly c. 750–530 Ma was produced via two main orogenies: the younger Kuunga Orogeny (c. 570–530 Ma) and the older is the EAO (~750–620 Ma) (Meert 2003).

The EAO extends for more than 6000 km in Gondwana. The preserved relics of the EAO include exhumed high-PT metamorphic belts of the Mozambique Belt (Southern India, Madagascar, East Africa) to greenschist or lower grade facies belts in the ANS (NE Africa and Arabia) (Cox et al. 2012). The EAO have been resulted from the collision of amalgamated arc terranes of the ANS with the Azania and Afif terranes to the east and the Sahara and Congo–Tanzania Cratons to the west (Fritz et al. 2013) (Figs. 2.11 and 2.12). Azania is a continental block between the Indian Shield and Congo–Tanzania–Bangweulu Craton, named after the classical name for the East African coast (Collins and Pisarevsky 2005). Also, Azania is defined as microcontinent of Archean and Paleoproterozoic crust (2900–2450 Ma) extend over Somalia, Madagascar and Arabia (Afif terrane) (Fritz et al. 2013) (Fig. 2.11). A Neoarchean continental block named Al-Mahfid Block of Yemen links between Azania and Afif (Windley et al. 1996; Whitehouse et al. 2001). As a result of the Mozambique Ocean rollback, the Azania Microcontinent detached from the Congo–Tanzania–Bangweulu Craton and the subducted oceanic crust beneath central Madagascar from a subduction zone marked by the eastern Betsimisaraka Malagasy Suture (Collins and Windley 2002; Collins et al. 2000; 2007 and 2012). Separation of Azania Microcontinent from the Congo–Tanzania–Bangweulu Block (~800–750 Ma) produced a back-arc basin that formed behind the continental arc created on Azania (Collins and Pisarevsky 2005). Both the eastern and western margins of Azania are marked by Neoproterozoic volcanosedimentary sequences and rocks formed in an oceanic environment (Collins and Pisarevsky 2005). The Azanian western margin of Kenya, Ethiopia, Eritrea, and Sudan, Egypt and Saudi Arabia is marked by ophiolites of Neoproterozoic age and juvenile volcanics (Abdelsalam and Stern 1996; Rogers and Santosh 2004). The ocean between Azania and East Africa was consumed in the Tonian along an intra-oceanic arc (Collins et al. 2012). The ocean closed at ~630 Ma by collision between the Azanian and Congo-Tanzanian continents (Collins and Windley 2002; Collins 2006). This collision produced highly grade metamorphosed and deformed rocks in eastern Africa (Hauzenberger et al. 2004, 2007) and southwestern Madagascar (Jöns and Schenk 2011). At the same moment, or not long after the East African Orogeny, a sequence of Cadomian terranes such as Avalonia and Armorica were rifted off the northern Gondwana margin early in the Paleozoic and accreted to Laurentia (Keppie et al. 2003; Fritz et al. 2012, 2013). The rifting of these terranes, likely due to the withdrawal of the Cadomian arc, may have enabled the ANS to be expelled to the north (Jacobs and Thomas 2004; Meert and Lieberman 2008; Fritz et al. 2012, 2013).

2.2.4 Kuunga Orogeny

The EAO resulted from the amalgamation of arc terranes in the northern ANS and continental collision between East African pieces and parts of the Azania terrane in the south (Collins and Pisarevsky 2005). The change from arc suturing to continental collision settings is seen in southern Kenya where southernmost arcs of the ANS merge with thickened continental margin suites of the Eastern Granulite Belt. The younger ca. 570–530 Ma Kuunga Orogeny heads from the Damara–Zambesi–Irumide Belts (De Waele et al. 2006) over Tanzania–Mozambique to southern India and clashes with the EAO in southern-central Tanzania (Fig. 2.12). Two transitional orogen settings may be defined, (1) that between island arcs and inverted passive continental margin within the EAO and (2) that between N–S trending East African and W–E trending Kuunga Orogenies. SE-Kenya’s Neoproterozoic arc suites are revealed as a thin stripe between west Azania and the eastern granulite belt. This suture is a long, NNW extended belt that roughly aligns with the prominent southern ANS shear zones that intersect at the ANS’s southern tip (Athi and Aswa Shear Zones) (Collins and Pisarevsky 2005; Fritz et al. 2009, 2012, 2013).

The Kuunga Orogeny (Fig. 2.12) is an orogeny that occurred during the Ediacaran and Cambrian in South-East Africa. It is slightly younger than the East African orogeny and produced by collisions between India and Australia–East Antarctica and Azania and India (Meert 2003; Collins 2003; Jacobs and Thomas 2004; Boger and Miller 2004; Collins and Pisarevsky 2005; Kelsey et al. 2008; Cox et al. 2012). It documents the collision between north and south Gondwana, or what is today Dronning Maud Land in Antarctica and northern Mozambique in Africa (Meert 2003; Grantham et al. 2013). India collided with Australia–East Antarctica and India collided with Azania in the Kuunga orogeny before the formation of Gondwana (Collins and Pisarevsky 2005; Kelsey et al. 2008; Cox et al. 2012). The proposed location of the Azania–India collision named the Malagasy Orogeny by Collins and Pisarevsky (2005) for the Ediacaran to Cambrian orogeny that resulted as India collided with Azania and the Congo–Tanzania–Bangweulu Blocks. The Kuunga Orogeny is prevalent in the southern East African orogeny in southern India (Fig. 2.12), Madagascar and central Arabia (Stern 1994a, b; Fritz et al. 2013). In East Africa, an E–W trending belt with Kuunga ages spreads from Zambia–Malawi across Mozambique and further into southern India and Sri Lanka (Collins et al. 2007; Plavsa et al. 2012; Fritz et al. 2013). Overprinting of Kuunga structures with the 620 Ma East African fabrics has led to complex patterns of interference in the Eastern Granulites.

2.3 The Arabian–Nubian Shield (ANS)

The ANS is southward narrowing belt that forms the suture shared between East and West Gondwana at the northern end of the EAO (Stern and Kröner 1993; Stern 2002; Stern et al. 2010; Ali et al. 2009a, b, 2012a, b; Johnson et al. 2011; Merdith et al. 2017). It stretches over 3500 km to the north and 1500 km to the east and west (Fig. 2.13), underlying an area of ~2.7 x 10⁶ km² in the northern half of the EAO (Stern 1994a, b; Johnson 2014). It is representative of the continental crust that underlies NE Africa, SW Asia and Arabia where the eastern half of the shield is a segment of the Arabian Plate, while the western part pertains to the African Plate (Johnson and Woldehaimanot 2003; Rogers and Santosh 2004).

The East Gondwana assembly was driven mainly by the Mozambique subduction recorded in the formation of two main orogens, Eastern Orogen and Kuunga (Meert 2003; Collins and Pisarevsky 2005; Merdith et al. 2017). West Gondwana is represented by Archaean and Palaeoproterozoic continental crust belonging to the Saharan Metacraton, which was strongly reworked by Neoproterozoic thermal and deformational events and is discontinuously exposed west of the Nile (Fritz et al. 2013). The Saharan Metacraton is a large area of continental crust in the north-central part of Africa. Other names have been used to describe the general area that reflects different views of its nature and extent. These include “Nile Craton”, “ Ghost Sahara Congo Craton” and “Eastern Saharan Craton” (Stern 2002). Abdelsalam et al. (2011) defined a metacraton as “a craton that has been remobilized during an orogenic event but is still recognizable dominantly through its rheological, geochronological and isotopic characteristics”. Liégeois et al. (2013) summarized the main distinctive characteristics of the metacraton as follows: (1) absence of pre-collisional events; (2) absence of lithospheric thickening, high-pressure metamorphism being generated by subduction, leading to high gradient in strain and metamorphic intensity; (3) preservation of pre-collision allochthonous ocean terranes; (4) abundant post-collisional magmatism associated with shear zones but not with lithospheric thickening; (5) presence of high-temperature–low-pressure metamorphism linked to post-collisional magmatism; (6) intracontinental orogenic belts not related to either subduction or oceanic basin closure.

2.3.1 Geology of the Arabian–Nubian Shield

The ANS consists of Precambrian metamorphic rocks cropping out along the coastline of the Red Sea, and exposed in areas of the Arabian and Sahara Deserts (Fig. 2.14), and in the highlands of Yemen, the Asir Province of Arabia and Ethiopian Highlands of the south (Figs. 2.11, 2.13 and 2.14). The ANS includes areas of western Arabia and northeastern Africa (Somalia, Yemen, Ethiopia, Eritrea, Sudan, Saudi Arabia, Egypt, Jordan and Palestine) (Johnson and Woldehaimanot 2003; Johnson et al. 2011; Fritz et al. 2012, 2013; Johnson 2014). Formation of arcs in the ANS occurred over a ~300-million-year period including supercontinent Rodinia break-up and the assembly of supercontinent Gondwana. The arc assemblages include metavolcanics and volcanoclastic metasediments as well as large amounts of diorite, tonalite, trondhjemite and granodiorite. Geochemically, the arcs have calc-alkaline to tholeiitic and locally MORB chemistry (Johnson 2014). The oceanic crust in the ANS is represented by ophiolite complexes, ranging from 845 to 675 Ma (Stern et al. 2004; Johnson 2014) (Fig. 2.14). The arc assemblages range from ~870 to 615 Ma (Stern et al. 2004). Based on differences in age of formation, structural style and, locally, isotopic characteristics, the arc assemblages are divided into tectonostratigraphic terranes (Johnson et al. 2011; Stern et al. 2004; Johnson 2014) (Figs. 2.11 and 2.13). Boundaries between the terranes are ophiolite-decorated shear zones or sutures and transcurrent shear zones (Johnson 2014) (Fig. 2.13). Most protoliths are early to middle Cryogenian, but Tonian rocks are locally preserved in the central and southern ANS (Johnson 2014). Arc amalgamation and suturing occurred between ~780 and 600 Ma. The accretion between the ANS and the Saharan Metacraton reflects a terminal collision between the ANS and West Gondwana blocks at approximately 650–580 Ma (Johnson et al. 2011; Fritz et al. 2012, 2013; Johnson 2014). Metamorphic grades in the ANS range from granulite facies in the south, decreasing to greenschist facies to the north and occurred during transpressional east–west shortening, north–south extension, strike–slip shearing and tectonic escape (Johnson 2014).

In the south, the west margin of the ANS is defined by juxtaposition of ophiolite and ophiolitic mélanges and juvenile Neoproterozoic arc magmatic terranes with the MB, the Sahara Metacraton and the Archean Congo–Tanzania Craton (Fritz et al. 2012, 2013) (Figs. 2.11, 2.13 and 2.15). The Congo Craton is an ancient Precambrian craton that with the Kaapvaal, Zimbabwe, Tanzania and West African Cratons form the modern continent of Africa (Rogers and Santosh 2004; Ernst et al. 2013). These cratons were formed between about 3.6 and 2.0 Ga and have been tectonically stable since that time (Rogers and Santosh 2004; Ernst et al. 2013). The Congo Craton occupies a significant portion of Central South Africa, from the Kasai region of the Democratic Republic of Congo to Sudan and Angola. It forms parts of the countries of Gabon, Cameroon and the Central African Republic. A small portion extends into Zambia as well, where it is called the Bangweulu Block (Ernst et al. 2013).

Most of the ANS western margin is not defined in the north, as it is covered by Mesozoic to Cenozoic sedimentary rocks (Fritz et al. 2013). In Sudan and Kenya, the western margin crops out in the Keraf Arc-Continent Suture (Abdelsalam et al. 1998), the Kabus ophiolite (Shellnutt et al. 2017, 2018), the Sekerr ophiolite of northwestern Kenya (Mosley 1993; Shellnutt et al. 2017) and the Kinyiki ophiolite (Fig. 2.13) of southern Kenya (Frisch and Pohl 1986). The Keraf Suture is an N-trending suture about 500 km long and ~50 km wide. It formed during Gondwana amalgamation and lies between the ANS in the east and the older Nile Craton to the west (Abdelsalam et al. 1998). The southern part of the suture is dominated by N- and NNW-striking, sinistral strike–slip shear zones, whereas the northern part is deformed by N-trending upright folds (Abdelsalam et al. 1998). The southern tip of the ANS is represented by old amphibolites and gneisses (955–845 Ma) close to Galana–Athi shear zone (Fig. 2.15) which is considered as part of Azania craton (Bauernhofer et al. 2008) (Figs. 2.11 and 2.13).

The eastern margin of the ANS in Kenya, Ethiopia and Somalia is marked by the Mutito-Bruna Shear Zone (Mosley 1993). Mutito-Bruna Shear Zone (Figs. 2.11 and 2.15) occurs in contact with Azania and represents intense ductile shearing producing closely spaced steep, upright, N–S or NNW–SSE ductile zones, often with a constrictional, linear tectonite fabric generally with a sinistral shear sense (Mosley 1993). The Mutito shear zone (Figs. 2.11 and 2.15) separates the Barsaloi-Adola Moyale ophiolitic belts from the poorly exposed older suites of the Burr Complex in southern Somalia (Mosley 1993). The boundary between shared by Azania and the ANS in Yemen is probably defined by arc-continent sutures along the Abas and Al-Mahfid terranes (Figs. 2.11 and 2.13). The boundaries between the ANS and the Afif Terrane and Khida Subterrane in eastern Saudi Arabia are arc–arc sutures (Johnson et al. 2011); therefore, they considered as crustal blocks within the ANS (Fig. 2.13).

The ANS initiated by at c. 870 Ma by rifting of Rodinia, closure of the Mozambique Ocean and accumulation of volcanic arcs during a few hundred million years before ~600 Ma (Rogers and Santosh 2004; Johnson et al. 2011; Fritz et al. 2013) on thin juvenile crust of the Mozambique Ocean. This region of juvenile crust extends to the northern margin of the Ethiopian flood basalts and also appears in scattered outcrops just south of the basalts (Rogers and Santosh 2004; Fritz et al. 2013). The ANS incorporates Middle Cryogenian–Ediacaran (790–560 Ma) terrestrial and shallow-marine sedimentary and volcanic successions (Fig. 2.16) which lie unconformably on Cryogenian juvenile crust. The older units were deposited in the central ANS after 780–760 Ma shearing and suturing. In the northern ANS, ~700 Ma suturing is consistent with Middle Cryogenic Basins. In the eastern ANS, Late Cryogenian basins were superimposed and accompanied by Nabitah Orogeny at 680–640 Ma (Johnson et al. 2011, 2013). Ediacaran successions are found in pull-apart and transpressive basins formed during final consolidation of the ANS (Johnson et al. 2011, 2013).

Ediacaran post-amalgamation basins are widespread in the ANS (Fig. 2.17). The fill of these basins includes Dokhan volcanics and sedimentary rocks belonging to Hammamat and Thalbah Groups, and the Saramuj Conglomerates (Johnson et al. 2011, 2013). Deposition of oceanic volcanic arcs was followed by intrusion of post-tectonic granites and stabilization of the terrane between 600 and 500 Ma, terminated at c. 550 Ma by the transformation of the northerly sections of the EAO into a passive margin on the southern shore of palaeo-Tethys (Fritz et al. 2013). The Kuunga shortening phase occurs under central Arabia’s Phanerozoic and is absent from the exposed ANS (Johnson et al. 2011; Cox et al. 2012).

2.3.2 Ophiolites of the Arabian–Nubian Shield

The ANS formed from ~870 to 550 Ma is one of the largest tracts of juvenile Neoproterozoic crust (Johnson 2014). The main diagnostic features of the ANS ophiolites are (i) greenschist facies metamorphism, (ii) dismembered, altered and consist mainly of tectonized harzburgites, pillowed basalts and gabbros as well as plagiogranite (Fig. 2.15), (iii) several ophiolites have sheeted dike complexes (such as Ghadir, Onib and Ess), (iv) shearing along NW-trending strike–slip faults and shear zones pertaining to the Najd Fault System especially ophiolites in the northern ANS (Sultan et al. 1988; Stern et al. 2004) and (v) form nappe complexes along terrane sutures, especially In the northern ANS.

Mid-Neoproterozoic ophiolites (690–890 Ma) are well represented in the ANS and extend 3000 km N–S and >1000 km E–W from the farthest north (Jebel Ess) almost to the equator, and from Rahib in the west to Jabal Uwayjah (45°E) in the east (Stern et al. 2004), encompassing an area of around two million square kilometres (Fig. 2.1, Stern et al. 2004). The ophiolite belts represent sutures marking the location of island arcs extending to those in Saudi Arabia on a pre-Red Sea drift reconstruction and further south to Mozambique (Fig. 2.15). North of the Bi’r Umq-Nakasib suture, the ophiolite belts strike nearly E–W (Stern et al. 2004). Bi’r Umq-Nakasib suture extends SW from Thurwah and Bi’r Umq Ophiolites in Arabia and extends over the Red Sea through Oshib and Meritri Ophiolites in Sudan.

Reconstruction of relative ages across the ANS demonstrates that the ophiolite belts young to the east (Berhe 1990). Geological and geochemical studies suggest rifting at c. 1200 Ma, and subsequent convergence led to the development of intra-oceanic arcs and associated marginal basins in the north and narrow basins further south in Kenya and Tanzania (Berhe 1990). Many of the ANS mafic–ultramafic complexes evolved in a seafloor spreading environment, while others are the substrate of island arcs (Quick 1990, 1991; Stern et al. 2004), while others may be layered intrusions (Stern et al. 2004). The Yubdo complex in Western Ethiopia consists of harzburgite, cumulate sequence of ultramafic, gabbroic rocks and metabasalts (Berhe 1990). The Baragoi complex in Kenya is composed of tectonized harzburgite, dunite with chromite pods, ultramafics and gabbroic rocks. Available evidence for the Adola Moyale Belt, in Southern Ethiopia and NE Kenya, point to MORB and island-arc geochemistry, of a back-arc environment (Berhe 1990). The Ingessana Complex in Sudan has an island-arc tholeiitic tectonic affinity that indicates development in a supra-subduction setting. The data from Sol Hamed Complex in Egypt suggest a back-arc basin (Berhe 1990).

ANS ophiolitic lavas often describe a subalkaline suite of low K and high Ti contents. They reveal both tholeiitic and calc-alkaline affinities and include a significant subordinate proportion of boninites Arabian shield (Stern et al. 2004). ANS ophiolites are frequently associated with a 1- to 3-km-thick sequence of ultramafic cumulates (dunite- and pyroxene-rich lithologies) transitioning into higher layered gabbroic rocks (Stern et al. 2004). A supra-subduction zone setting for the Egyptian Eastern Desert ophiolites is commonly accepted (e.g. El Bahariya and Arai 2003; Azer and Khalil 2005; Azer and Stern 2007; Farahat et al. 2011; Ahmed et al. 2012; Abdel-Karim et al. 2016; El Bahariya 2018). Furthermore, a back-arc setting is generally favoured (El-Sayed et al. 1999; Ahmed et al. 2001; El Bahariya and Arai 2003; Farahat et al. 2004; El Gaby 2005; Abdel-Karim et al. 2008; El Bahariya 2018). Also, a fore-arc environment was proposed on the basis of whole-rock chemistry and mineral analyses of serpentinites (Stern 2004; Azer and Stern 2007; Azer et al. 2013).

Neoproterozoic ophiolites and ophiolitic mélange (El Bahariya 2012) are widespread in the southern and central sectors of the ENS. The components of ophiolites are dismembered and altered and include harzburgite, cumulate ultramafics, sheeted dikes, pillowed basalt, layered gabbro and plagiogranite (Stern et al. 2004). A complete ophiolite section was detected in a few places in the ENS (e.g. Gerf, Fawkhir, Ghadir) including serpentinized ultramafics, layered and isotropic gabbros, sheeted dykes and mafic lavas including pillow basalts. Mostly, the Egyptian ophiolites occur as tectonized masses, sheets, lenses and mélanges of these ophiolitic components (El Sharkawy and El Bayoumi 1979; Azer et al. 2013). The ophiolites in the Central Eastern Desert show characteristics of both MORB-type and SSZ-type ophiolites and classified from oldest to youngest by El Bahariya (2018) based on mode of occurrences and geological characteristics into (1) MORB intact ophiolites (e.g. Muweilih, Ghadir), (2) dismembered ophiolites and (3) arc-associated ophiolites (e.g. El Sid, Esel). Dismembered ophiolites occur either as (i) ophiolite blocks in mélanges (e.g. Kareim El Abiad tectonic mélange, Garf tectonic mélange, Muweilih olistostrome, Esel olistostrome, Um Esh olistostromal mélange and Mubarak olistostromal mélange) or as (ii) ophiolites along structural contacts (Abu Meriewa occurrence, Sodmien occurrence, Um Saneyat occurrence and Um Khariga occurrence (El Bahariya 2018).

2.3.3 How Juvenile Is the Arabian–Nubian Shield?

As mentioned before, the ANS is the northern half of a great collision zone called the EAO. This collision zone was created at the end of the Neoproterozoic time, when East and West Gondwana combined to create the Gondwana supercontinent. Across what is now Southern Africa, the most intense part of the collision occurred, where older crust in Tanzania, Mozambique and Madagascar was remobilized to form the Mozambique Belt (MB). This great collision was responsible for the Pan-African Orogeny near the end of Neoproterozoic time (Kröner and Stern 2004). The crust of the MB is quite different from that of the Arabian–Nubian Shield, which is predominantly “juvenile” crust, that is, crust that formed from partial melting of Earth’s mantle, although much older Archean and Paleoproterozoic crustal materials, is exposed west of the Nile in Egypt, in the SE part of the shield in Arabia, in eastern Ethiopia and in Yemen (Kröner and Stern 2004).

The Arabian shield retains a sustained magmatic record of amalgamated juvenile terranes that contain a wide variety of early Neoproterozoic to Cambrian granitoids intruding into volcanosedimentary basin assemblages. Robinson et al. (2014) provided U–Pb zircon dating of 19 examples of granitoids that intruded eight terranes of the Arabian shield and defined four separate magmatic events: anorogenic (∼525 Ma), post-tectonic (∼620 Ma), syncollisional (∼710 Ma) and island arc (∼845 Ma).

With the development of island-arc magmatism associated with the closing of the Mozambique Ocean, syncollisional magmatism in the ANS at 715 Ma reflects the first level of terrane amalgamation in the Mozambique Ocean due to the convergence associated with East Africa and East Gondwana subduction. The accretion/arc collision in the western ANS (west of Afif) was terminated at ∼630 Ma due to collision of Azania with the Congo–Tanzania–Bangweulu Block (Collins and Windley 2002; Collins and Pisarevsky 2005). The emergence of post-tectonic magmatism intruding the Nabitah Belt and the creation of western ANS back-arcs reinforce this understanding. East of Afif, the appearance of ∼611–607 Ma post-tectonic magmatism cross-cutting the Halaban Suture, possibly corresponds with the collision of the Afif-Abas Block and eastern Arabia (Collins and Pisarevsky 2005). The final phases of the Gondwana assembly are reported from ~5570–530 Ma in the northern EAO and split into two orogenic periods, the Kuunga and Malagasy, respectively (Collins et al. 2014 and references therein). It is proposed that India’s interaction with the Congo–Tanzania–Bangweulu Block (Malagasy Orogeny) has produced the ca. 525 Ma A-type magmatism (Hijaz Terrane). It is well known that the major transform fault systems such as the NFS are spread throughout the northern Arabian shield and shape a sequence of reactivated events that cause magmatism (Stern 1985; Johnson et al. 2011) such as ∼525 Ma Mardabah Complex.

However, no crust older than ∼890 Ma (Deschamps et al. 2004; Hargrove et al. 2006) is exposed in the core of the juvenile ANS, and crust of Mesoproterozoic age is not exposed anywhere in the ANS. Hargrove et al. (2006) concluded that the ANS is less juvenile than presently appreciated and may contain a significant amount of continental crust and estimates of crust-formation rates that cite the ANS as a model for juvenile crust should consider this possibility.

2.3.4 Accretion and Thickening in the Arabian–Nubian Shield

The relics of the EAO (≈6000 km long in Gondwana) include high-grade metamorphic complexes in the Mozambique Belt, and greenschist or lower grade belts of volcanics, plutons and sedimentary rocks in the Arabian–Nubian shield (Cox et al. 2012). The metamorphism connected to the closing of the Mozambique Ocean in southern India and eastern Madagascar was linked to an Ediacaran collision between the Congo–Tanzania–Bangweulu Block and India in the Neoproterozoic time (Clark et al. 2009; Collins et al. 2003, 2007; Cox et al. 2012) that had previously accreted a continent known as Azania during the Cryogenian (Collins 2006; Collins and Windley 2002) to form the Malagasy Orogen (Collins and Pisarevsky 2005).

The ANS took about 300 million years to develop. The oldest rocks related to the ANS crust’s formative cycle created by the coalescence of island-arc and back-arc basins and even oceanic plateaus. The oldest rocks in this series are around 870 Ma and are located in E Sudan and southeast Saudi Arabia (Whitehouse et al. 2001; Collins and Windley 2002; Hargrove et al. 2006). The oldest units are commonly ophiolites, which demonstrate that the creation of continental ANS crust began with the forming of oceanic crust by the expansion of the seafloor, accompanied by subduction and island-arc development (Johnson et al. 2011). The early magmatic stages in the ANS involve the presence of dense oceanic tholeiite sequences in intra-oceanic settings accompanied by bimodal volcanism of island-arc chemistry 950–650 Ma. Such processes culminated in huge amounts of mafic crust and lithospheric mantle within the ANS and crustal overprinting (Bentor 1985; Stein and Goldstein 1996; Stein 2003) for producing, these upwelling plumes possibly provide the lithophile element-enriched sources for younger calc-alkaline (640–590 Ma) and alkaline (590–550 Ma) magmatism (Robinson et al. 2014).

Over the time 780–620 Ma, various island arcs collided and these tectonic terranes sutured and developed a progressively large and dense nucleus of continental juvenile crust. This thickening resulted in the formation of several suture zones, marked by obduction of ophiolites and intense deformation (Johnson et al. 2011; Cox et al. 2012). Crustal thickening was also accompanied by melting and magmatic fractionation of mafic magmas that ponded deep in the crust. These melts rose upwards to be emplaced as granitic plutons. Magmatism during this episode is characterized by tholeiites and calc-alkaline suites (Robinson et al. 2014).

The juvenile ANS crust was trapped between great tracts of converging continental crust. A protracted episode of continental collision started at about 610 Ma ago and continued for about 50 million years. Collision was more intense in the south, but it also strongly affected the ANS (Kröner and Stern 2004). The arc terranes and sutures in the southern ANS were deformed by N–S-oriented upright folds and shear zones forming elongated systems such as the Hamisana Shear Zone in North Sudan and South Egypt (Johnson and Woldehaimanot 2003). The ANS was influenced more far north and east by the development of the broad NW-SE pattern of the Najd Fault System (NFS). The composition of igneous rocks became distinctively more evolved as the collision continued and the crust continued to thicken. Deep erosion, possibly by a continental ice sheet, happened during this time. All tectonic and magmatic activities ended by the time the Cambrian sandstones were deposited, about 530 million years ago (Stein 2003; Robinson et al. 2014). A variety of characteristics have been traced to late stage extension tectonics, including a large NE–SW trending dyke swarm, NE–SW trend normal faults and NW-SE trend sedimentary basins filled with post-orogenic molasse deposits (Blasband et al. 2000; Johnson and Woldehaimanot 2003; Kröner and Stern 2004).

2.3.5 Tectonic Models of the Arabian–Nubian Shield

The ANS contains a vast expanse of Neoproterozoic juvenile oceanic arc crust surrounded by older cratonic crust to the west and east (Stoeser and Frost 2006). Fritz et al. (2013) and Hamimi et al. (2019) recognized five main phases produced the ANS: (1) rifting of the African craton (∼1200–950 Ma); (2) ensimatic island-arc development (∼950–715 Ma); (3) formation of the Arabian–Nubian Neocraton by microplate accretion and continental collision (715–640 Ma); (4) collision-related intra-cratonic magmatism and tectonism (∼640–550 Ma); and (5) epicontinental subsidence (<550 Ma), commonly found in fault-bound basins (Nettle et al. 2013).

The Arabian shield in Western Saudi Arabia was developed in Neoproterozoic through three main stages: (1) Island-arc creation and accretion, primarily during the time ∼870–620 Ma; (2) Continental orogenesis after collision with East Gondwana’s north-western margin, ∼660–620 Ma; and (3) post-collisional extension, magmatism and sedimentation ∼620–540 Ma (Greenwood et al. 1976; Al-Shanti and Mitchell 1976; Schmidt et al. 1979; Stoeser and Camp 1985; Kröner 1985; Genna et al. 2002; Johnson and Woldehaimanot 2003; Stoeser and Frost 2006). The collision was occurred in the south, whereas subduction of oceanic crust continued in the north until approximately 620 Ma (Doebrich et al. 2007); therefore, there is overlap in ages for arc formation and collision.

During collisional orogenesis, the Najd Fault System (NFS) (a major left-lateral fault system) developed throughout the northern part of the Arabian shield. The Najd was described as the product of tectonic escape, similar to those arising from the collision between India and Asia (Schmidt et al. 1979; Burke and Sengor 1986). Three of the most prominent of the Najd Fault Zones, Ruwah, Ar Rika and Halaban-Zarghat (Al-Saleh et al. 1998) are shown in Figs. 2.17 and 2.18. Moore (1979) described this system as a major transcurrent (strike–slip) fault system, of Proterozoic age in the Arabian shield. The NFS was identified originally as a huge late Proterozoic and early Phanerozoic NW-trending brittle–ductile shear zone with 300 km width and length over 1100 km extending over the north part of the ANS (Stern 1985; Johnson et al. 2011). The displacement along the strike of the NFS was reported by Brown (1972) and Hamimi et al. (2019) as 240 km cumulative displacement but field displacements can be demonstrated as only tens of kilometres for faults (Johnson et al. 2011). The NFS and other NW-trending strike–slip faults in the ANS are known to be post-accretionary structures and were deduced from the squeezing of the Arabian–Nubian shield between East and West Gondwana (Berhe 1990; Stern 1994a, b; Abdelsalam 1994; Abdelsalam and Stern 1996; Abdelsalam et al. 2003; Hamimi et al. 2019). The formation of NFS is a result of simple shear that allowed the Nubian and southern Arabian shield to move several hundred kilometres sinistrally with respect to northern Arabia (Moore 1979; Hamimi et al. 2019).

Stoeser and Frost (2006) recognized two stages for terrane accretion of the Arabian shield: (1) Accretion of terrane emerging from the closure of ocean basins (e.g. Stoeser and Camp 1985; Johnson and Woldehaimanot 2003) and (2) Lateral transpressional accretion as suggested for the North American Western Margin (Samson and Patchett 1991). Furthermore, although most theories include mainly the accretion of oceanic island arcs, some have indicated that the accretion of oceanic plateaus has also played an important role in the evolution of the ANS (e.g. Stein and Goldstein 1996).

The terrane concept has been applied to the ANS since the mid-1980s, developing out of earlier plate-tectonic/subduction zone and accreted island-arc concepts, and leads to the interpretation that the exposed and immediately adjacent concealed Proterozoic rocks of western Arabia and northeastern Africa are a collage of crustal blocks, or terranes that, with varying degrees of confidence, can be correlated across the Red Sea, thereby linking the now separated Arabian and Nubian shields. Given the reconnaissance nature of geologic mapping in the Arabian shield, details about the origins, ages, structure and convergence of the terranes remain to be established, and the terrane-collage model is, strictly speaking, a working hypothesis.

2.4 The Egyptian Nubian Shield

Although Precambrian rocks underlie the entire Egypt and Sudan, they are best known where exposed in the Nubian Shield, in the Egyptian Eastern Desert and Northern Sudan. In the shield, the Precambrian rocks are mostly Neoproterozoic. The tectonic development of the ANS spans three phases spanning over 600 Ma: accumulation of arc terrains within the Hijaz Magmatic Arc, accompanied by accretion of the Hijaz Magmatic Arc towards the Nile Craton and reworking of the accreted arc after accretion (Fritz et al. 1996; Abdelsalam and Stern 1996; Augland et al. 2012). The Egyptian Nubian Shield (ENS) covers ~100,000 km², crops out along the Red Sea Hills in the Eastern Desert and southern Sinai, as well as limited areas in the south Western Desert (Oweinat area) (El Gaby et al. 1990; Hassan and Hashad 1990). The basement complex of the ENS comprises four main tectonostratigraphic units (Hamimi et al. 2019) from bottom upward as follows: (1) high-grade gneisses and migmatites, (2) arc-type volcanic/volcanosedimentary units, along with dismembered ophiolites, (3) the Ediacaran Hammamat and Dokhan supracrustal sequences, and (4) granitoids. During the Cambrian until the Cretaceous, the Nubian Shield’s rocks were almost constantly hidden by Phanerozoic sedimentary rocks, until the uplift and deforestation correlated with the division of Arabia and Africa culminated in the Nubian shield becoming revealed as it is today. Juvenile Neoproterozoic rocks predominate in the Nubian shield and extend distances of 1000 km from northern Sudan to northern Egypt.

Four major lithological components describe the ENS (Fig. 2.19): predominantly juvenile arc, supracrustal series, ophiolites and gneiss domains enclosing core complexes and intrusions of granitoids (Fritz et al. 1996, 2002; Abd El-Naby et al. 2000, Shalaby et al. 2005; Abd El-Wahed 2008, 2014; Ali et al. 2015; Hamimi et al. 2019; Abd El-Wahed et al. 2019a, b; Hamimi et al. 2019; Zoheir et al. 2019b; Hamimi and Abd El-Wahed 2020; Fowler and Hamimi 2020; Abd El-Wahed and Hamimi 2020). These units were tectonically mixed by thrusting associated with accretion and sinistral strike–slip shearing along the Najd and other NW-striking shear zones, particularly in the central part of Eastern Desert of Egypt (Abd El-Wahed 2010; Hamimi et al. 2019). Sheared granitoid gneisses occur locally in antiformal “domal” locations in the Eastern Desert of Egypt, typically surrounded by low-grade supracustal assemblages, such as the Meatiq (Sturchio et al. 1983; Fritz et al. 1996; Andresen et al. 2009, 2010; Hamdy et al. 2017; Mohammad et al. 2019), Hafafit (Greiling et al. 1994, Abd El-Naby et al. 2008, Lundmark et al. 2012; Ali et al. 2015; Makroum 2017; Hamimi et al. 2019), El-Shalul (Hamimi et al. 1994, Osman 1996, Ali et al. 2012a) and in the region along Wadi Beitan (Abdel Khalek et al. 1992; Ali et al. 2015). The gneiss domes of the Egyptian Eastern Desert were previously interpreted to represent a pre-Neoproterozoic (pre-Pan-African) basement (e.g. El-Ramly and Akaad 1960; El-Ramly 1972; El-Gaby et al. 1984, 1988, 1990; Hamimi et al. 1994). However, other authors argued that the Eastern Desert gneissic rocks are of juvenile origins and formed in intra-oceanic or continental arc settings within, or on the margin of, the Mozambique Ocean, or along one or more magmatic arcs along the western margin of this ocean prior to amalgamation of East and West Gondwana at ∼630 Ma (El-Ramly et al. 1984, 1993; Greiling et al. 1984, 1988; Stern 1994a, b; Abd El-Wahed 2008, 2014; Andresen et al. 2009, 2010; Abd El-Wahed and Kamh 2010; Augland et al. 2012, Lundmark et al. 2012; Ali et al. 2009a, b; 2010a, b, 2012a, b, 2015; Abd El-Wahed et al. 2016, 2019a, b; Zoheir et al. 2019b: Hamimi et al. 2019; Hamimi and Abd El-Wahed 2020; Fowler and Hamimi 2020; Abd El-Wahed and Hamimi 2020).

The Egyptian Eastern Desert was subdivided by Stern and Hedge (1985) into Northern, Central and South-Eastern Desert regions (NED, CED and SED); all reveal different aspects of the region’s protracted and intense Neoproterozoic episode of deformation and igneous activity.

2.4.1 The South-Eastern Desert (SED)

The details of the Neoproterozoic of the SED are still poorly known. The SED differs from the CED in several styles (Stern 2018): (i) it is much less studied; (ii) it lacks interesting sediments such as BIF and diamictite; (iii) it has some volcanogenic massive sulphides, whereas the CED does not (Abd El-Rahman et al. 2012); (iv) there are no major sedimentary or volcanic Ediacaran successions as described in the NED and CED, such as the Hammamat Group and Dokhan Volcanics; and (v) it contains equal amounts of serpentinites to those of CED except for Abu Dahr ophiolites (Gahlan et al. 2015) and those defining the Allaqi–Heiani Suture in the southern SED (Zimmer et al. 1995; Gahlan and Arai 2009; Ali et al. 2010a; Azer et al. 2013). The SED usually tends to be more deeply exposed than the CED and is less influenced by Najd shearing (Stern 2018). Gneisses in the ENS represent the infrastructure rocks that exposed in tectonic windows beneath the suprastructure rocks. Gneisses and related rocks in the SED are sheared and ductility-deformed at shallow depths such as the gneisses of El-Hudi, Haimur-Abu Swayel, Abu Fas, Beitan, Khuda, Sikait and Hamisana.

The SED is characterized by the presence of Allaqi–Heiani Suture (AHS) and the Hamisana shear zone (HSZ). The AHS is a part of the megascopic Allaqi–Heiani–Oneib–Sol Hamed–Yanbu Suture (AHOSHYS) (Abdelsalam and Stern 1996, Abdelsalam et al. 2003; Zoheir et al. 2019a) that extends ∼E–W for at least 600 km, from the western edge of the shield in Egypt to the eastern edge of basement exposures in NW Arabia (Ali et al. 2010a). The AHS tracks eastward from Wadi Allaqi’s entry into Lake Nasser along the Egyptian–Sudan border to Gabal Gerf along the Red Sea (Ali et al. 2010a). Here, the suture is thought to be displaced southwards along the N–S HSZ and continues NE along the Onib segment to the Red Sea, across which it can be traced into northwestern Saudi Arabia (Ali et al. 2010a). The Allaqi–Heiani belt is an arc–arc suture accreted and formed between ∼730 and 709 Ma (Ali et al. 2010a, b). It separates the Midyan–South-Eastern Desert terrane in the North from the ∼830–720 Ma Hijaz–Gebeit terrane to the South (Shackleton 1994; Abdelsalam and Stern 1996; Abdelsalam et al. 2003; Hamimi et al. 2019; Zoheir et al. 2019a). In the central and western portions of the Allaqi–Heiani suture, ophiolites form imbricate thrust-bound sheets and slices of serpentinite, amphibolite, metagabbro and metabasalt, embedded in a strongly tectonized carbonate talc matrix, creating an elongated WNW–ESE or E–W belt (Abdelsalam and Stern 1996; Abdelsalam et al. 2003; Zoheir and Klemm 2007).

The HSZ is an ophiolite-decorated post-accretionary structure (Stern et al. 1990) and not a suture as suggested by Berhe (1990). The AHOSHYS is broad and consists of four major lithologic associations: (i) ophiolites, (ii) island-arc metavolcanic and metasedimentary successions, (iii) gabbroic to granitic intrusions and (iv) gneiss (Stern et al. 1990, Abdelsalam et al. 2003, Zoheir and Klemm 2007; Gahlan and Arai 2009; Ali et al. 2010a). The Allaqi segment of the AHOSHYS separates the SE segment of the Eastern Desert Terrane in the north from the Gabgaba Terrane to the south (Abdelsalam and Stern 1996; Ali et al. 2010a).

2.4.2 The Central Eastern Desert (CED)

The CED is in several respects the best known of the three subdivisions of the Eastern Desert, as its supracustal sequences are especially fascinating and insightful. The CED was characterized by two prominent tectonostratigraphic units; (i) The lower unit is commonly referred to as the structural basement (Infrastructure) (El-Gaby et al. 1990) or as ‘‘lower tier” (e.g. Bennett and Mosley 1987; Greiling et al. 1994) and comprises high-grade metamorphic gneisses, migmatites, schists and amphibolites; and (ii) the second unit is commonly referred to as structural cover (Suprastructure) or the Pan-African nappes and is described as the upper unit including low-grade metamorphosed ophiolite slices (serpentinites, pillow lavas, metagabbros), arc metavolcanics and arc metasediments (e.g. El-Gaby et al. 1990; Gahlan et al. 2012). Abdel-Karim and Ahmed (2010) summarized what is known about 38 different occurrences of ophiolitics in CED and SED. Greenschist facies arc volcanics and volcanoclastic metasedimentary rocks are important supracrustal constituents of the CED above the ophiolites (Figs. 2.20 and 2.21). Collectively, the infrastructure and supracrustal rocks were intruded by syn-tectonic calc-alkaline granites and metagabbros–diorite complex and covered by Phanerozoic sediments (Abd El-Wahed et al. 2010).

The ensimatic (oceanic) assemblage in the CED represents the oldest units, with ∼750 Ma U-Pb zircon ages (Kröner et al. 1992; Andresen et al. 2009; Ali et al. 2009b; Stern 2018). It is perforated with Cryogenian I-type granitoids and Ediacaran A-type granites (Ali et al. 2012b) and dislocated by a number of magmatic–metamorphic core complexes (Stern 2018).

The later stages of the CED’s crustal growth are distinguished by eruption of Dokhan volcanics that is consistent with the deposition of molasse-type Hammamat molasse sediments formed in terrestrial alluvial fan/braided stream environment (Grothaus et al. 1979; Abd El-Wahed 2010; Bezenjania et al. 2014). In response to the differential relief between more elevated regions of the NED and topographically lower areas of the CED which followed Najd fault deformation of the CED and volcanic rifting in the NED, Hammamat basins developed (Stern 2018). In addition, the crustal rocks were intruded by late to post-tectonic granites.

The CED is mostly distinguished by the domination of NW-striking tectonic structure (Fig. 2.21) marking the NW–SE sinistral shear zone of the NFS (Abd El-Wahed and Kamh 2010; Abd El-Wahed 2014; Abdeen et al. 2014; Abd El-Wahed et al. 2016, 2019a, b; Hamimi et al. 2019; Hamimi and Abd El-Wahed 2020; Fowler and Hamimi 2020; Abd El-Wahed and Hamimi 2020). Nappe transport directions reported from the CED show a variance from top to the NE (e.g. El Bayoumi and Greiling 1984), top-to-the-NW (Ries et al. 1983; Greiling 1987), top-to-the-SE (Kamal El Din et al. 1992) and top-to-the-SW (Abdeen et al. 2002; Abdelsalam et al. 2003). Oblique convergence transpression (Abd El-Wahed and Kamh 2013), whether dextral or sinistral, is a current mechanism adopted by many authors (Fritz et al. 1996, 2002, 2013; Loizenbauer et al. 2001; Makroum 2001; Bregar et al. 2002; Helmy et al. 2004; Shalaby et al. 2005; Abd El-Wahed 2008, 2010; 2014; Shalaby 2010; Abd El-Wahed and Kamh 2010; Zoheir and Lehmann 2011; Zoheir and Weihed 2014; Abd El-Wahed et al. 2016; Abd El-Wahed and Thabet 2017; Makroum 2017; Hamdy et al. 2017; Stern 2018) to explain and imply the deformation styles in the CED.

Simply and collectively, Abd El-Wahed and Kamh (2010) listed the deformation events of the CED as follows: (1) D₁ associated with NNW-ward thrusts; (2) D₂ related to NE- and SW-directed thrusts; (3) D₃ attributed to sinistral movement along the NW-striking shears of the NFS; (4) D₄ associated with dextral movement along NE-trending shear zones; and (e) D₅ later events. The NNW-directed thrusts were created as a result of oblique convergence during island-arc accretion (690–650 Ma) in Sibai gneisses (Bregar et al. 2002; Fowler et al. 2007; Augland et al. 2012) and 630 and 609 Ma in Meatiq and Hafafit gneiss-cored domes (Andresen et al. 2009). The tectonic system changed from the compressional arc accretion setting to the sinistral transpressional regime between 650 and 540 Ma. In the CED, exhumation of the core complexes (e.g. Meatiq, Sibai, El-Shalul and Hafafit) is related to the sinistral slip along the NFS, which bound them from the SW and NE (Fritz et al. 1996). Not only core complexes of CED but also the Hammamat sediments are tectonically affected by the NFS (Abd El-Wahed 2010).

The NFS and the other NW-trending strike–slip shear zones of the ANS were initially post-accretionary structures and were deduced from the stretching of the ANS between East and West Gondwana (Berhe 1990; Stern 1994a, b; Abdelsalam and Stern 1996; Abdelsalam et al. 2003). The NFS comprises brittle and ductile shears in a region up to 300 km in width and about 1,100 km long, including the north part of the Arabian shield.

Different scenarios were suggested to interpret one of the major landmarks of the CED, the gneissic domes or core complexes (Meatiq, Sibai, El-Shalul and Hafafit). Gneiss domes have been diversely interpreted as (1) antiformal stacks produced by thrusting (Greiling et al. 1994), (2) core complexes associated with NW-SE crustal extension (Fritz et al. 1996; Bregar et al. 2002; Abd El-Wahed 2008; Makroum 2017; Hamdy et al. 2017; Abd El-Wahed et al. 2019a, b; Abd El-Wahed and Hamimi 2020) and (3) interference patterns of sheath folds (Fowler and El Kalioubi 2002—for the Hafafit Complex). Recently, there are two tectonic models to explain the formation and exhumation of these core complexes in the CED: (1) Orogen-parallel extension model (Fritz et al. 1996) assigned evolution of the core complexes to sinistral shear activity on NFS-related shear zones and (2) extensional tectonics model as a result of simultaneous NW–SE folding and NW–SE extension (Fowler et al. 2007; Andresen et al. 2010).

Nugrus shear zone is one of the CED’s main NW-trending shear zones that divides the CED from Egypt’s SED and forms the border between the Wadi Ghadir region (East) and the Hafafit Core Complex (West). Nugrus shear zone starts with a thickness of 750 m retaining its width at least as far as Gabal Sikait, where its thickness decreased by 50 m (Harraz and El-Sharkawy 2001; Fowler and Osman 2009). Lundmark et al. (2012) dated to ≤595 Ma the sinistral shearing on the Nugrus Shear Zone.

Nugrus shear zone is interpreted as thrust showing top-to-W- or SW tectonic transport over the continental margin (El Bayoumi and Greiling 1984; El-Ramly et al. 1984; El-Gaby et al. 1988) or as roof thrust linked up NW-dipping thrust imbricates of gneissic rocks and allowing NW-ward displacement of low-grade CED metavolcanics over the gneisses (Greiling et al. 1994 and Greiling 1997). Many workers considered Nugrus shear zone as a Najd-related ductile strike–slip shear (Fritz et al.1996; Shalaby et al. 2005; Abd El-Wahed et al. 2016; Makroum 2017). Fowler and Osman (2009) described Nugrus shear zone as an example of E–W striking low-angle normal ductile shear with low-angle N-dip.

The NW-trending sinistral shear zone-related transpression is conjugated and overprinted by the dominant dextral transpression along NE–SW trending shear zones (Shalaby et al. 2005; Abd El-Wahed and Kamh 2010) (Fig. 2.21). Abd El-Wahed (2014) documented a huge NE-trending shear belt (up to 110 km in length) occupying the area between Wadi Barramiya and Wadi Sha’it (Fig. 2.21) to the west (up to 60 km in width) and extends through the whole width of the Central Eastern Desert to include the area between Wadi Mubarak and Wadi Ghadir on the Red Sea coast (up to 120 km in width) (Fig. 2.21). The 40–60-km-wide, ENE–WSW trending Mubarak–Barramiya Shear Belt (MBSB) deforms supracrustal successions and structures associated with the NW-trending shear fabric (Fig. 2.21). The MBSB is interpreted as a dextral transpressive shear zone ceased the extension of the NSZ and modelled in terms of a regional “flower structure” (Abd El-Wahed and Kamh 2010).

2.4.3 The Northern Eastern Desert (NED)

The NED (Fig. 2.20) is remarkably different from both the CED and the SED. The NED is characterized mainly by (i) a wide range of Ediacaran igneous rocks and associated Hammamat sediments (Wilde and Youssef 2000), (ii) absence of ophiolites and scarcity of gneisses, (iii) BIF and diamictite is unknown, (iv) absence of Najd deformation, (v) abundance of dike swarms that trend E–W to NE–SW, (vi) abundance of epizonal A-type granites indicating strong extension, and (vii) abundance of Dokhan Volcanics (Stern and Hedge 1985). The Dokhan volcanics occupy a limited area of the Egyptian basement and are best exposed north of latitude 26°N in the Eastern Desert (e.g. NED and northernmost CED) particularly at their type locality Gabal Dokhan. Dokhan Volcanics are unmetamorphosed rocks felsic to intermediate composition and have medium- to high-K calc-alkaline affinities. Dokhan Volcanics have been formed in subaerial environment and commonly associated and/or interbedded with Hammamat molasse sediments (Abd El-Wahed 2010; Eliwa et al. 2014b; El-Bialy 2020). Breitkreuz et al. (2010) reported the eruption of the Dokhan Volcanics during two major volcanic pulses: 630–623 and 618–592 Ma. The Hammamat molasses of the ENS were developed between 650 and 580 Ma in individual basins in variable tectonic settings (Fritz and Messner 1999; Shalaby et al. 2006; Abd El-Wahed 2010). Dessouky et al. (2019) indicated that the Hammamat molasse sediments have been deposited during three tectonic settings: (i) pre-collision stage comprise fore‐arc, intra‐arc and back‐arc basins; (ii) syn‐collision stage include fore‐arc basins and low land basins; and (iii) late‐collision stage embrace low land and intramontane basins.

The most exciting recent developments in the studies of the NED are the presence of small tracts of Cryogenian and Tonian igneous rocks (Stern 2018) including the 666 Ma Mons Claudianus granodiorite (Stern 2018), Tonian muscovite tonalite and I-type muscovite-bearing trondhjemite and granodiorite with complexity in the U-Pb zircon ages at ∼740 Ma (Eliwa et al. (2014a), the presence of ∼720 Ma ignimbrites unconformably overlying the ∼740 Ma granitoids (Bühler et al. 2014), the presence of ∼780 Ma dacite (Abd El-Rahman et al. 2017).

2.4.4 The Egyptian Nubian Shield in Sinai

The basement rocks in Sinai (Fig. 2.19) comprise four complexes: the Feiran–Solaf, Sa’al–Zaghra, Kid and Taba Metamorphic complexes. The metamorphic complexes are composed of orthogneisses and metasedimentary–metavolcanic sequences. They are surrounded by younger non-metamorphosed intrusions and volcanosedimentary sequences and metamorphosed at greenschist to amphibolite and rarely up to granulite facies conditions (Shimron 1984; Jarrar et al. 1993; El-Shafei and Kusky 2003; Abu El-Enen et al. 2003, 2004; Fowler and Hassan 2008; Abu-Alam and Stüwe 2009, 2010; Azer and Farahat 2011).

These metamorphic complexes were considered to be pre-Pan-African by El-Gaby et al. (1990) and Abu Anbar et al. (2004). Nonetheless, other studies have recorded geochronological evidence that point to the Pan-African period (Bielski 1982; Stern and Manton 1987; Eliwa et al. 2008). A younger tectono-magmatic period in Sinai is distinguished by large intrusions of calc-alkaline magma during ca. 635–590 Ma (Ali et al. 2009a, b; Be’eri-Shlevin et al. 2009). Post-collisional calc-alkaline granitoids were placed from c. 635 to 590 Ma, while the alkaline granite suite was developed over a span of c. Ma 608–580 (Be’eri-Shlevin et al. 2009). Despite close association in space and time, calc-alkaline and alkaline granitoids are clearly distinguished in some major and trace element characteristics. This was demonstrated for 27 key plutons and plutonic complexes (Be’eri-Shlevin et al. 2009; Azer et al. 2010; Eyal et al. 2010; Azer and Farahat 2011). The alkaline activity was dominated by bimodal alkaline volcanism and shallow intrusions of granitoids, which are mostly A-types, along with dike swarms (Beyth et al. 1994; Kessel et al. 1998; Samuel et al. 2007; Eyal et al. 2010). The volcanosedimentary successions of southern Sinai (Kid, Ferani, Rutig, Sa’al–Zaghra, Iqna Shar’a, Khashabi and Meknas) have been associated with the Dokhan–Hammamat successions (Shimron 1980; Bentor 1985; El-Gaby et al. 1991; Basta 1997; Azzaz et al. 2000; Moussa 2003; Azer 2007; El-Bialy 2010; Azer and Farahat 2011).

The Katherina Volcanics at Gabal Ma’ain (580–590 Ma) in the Sinai consists of porphyritic rhyolite lavas (450 m thick) with minor pyroclastics intruded by alkaline granitoids (578 ± 8 Ma) (El-Bialy and Hassen 2012). The Tarr Albitites complex comprises intrusive and volcanic eruptions of albitite surrounded by an emplacement breccia containing veins and dykes of carbonates (Azer et al. 2008). Zircons separated from the albitites yield a U–Pb age of 605 ± 15 Ma, and this is also the age of the intrusive carbonates (Azer et al. 2008). Geochemical signatures of Tarr Albitites indicate that they represent pristine igneous rocks and are not metasomatic products (Azer et al. 2008)

2.5 Summary and Concluding Remarks

1.
Rodinia was created at c. 1.23 Ga by accretion and aggregation of fragments formed by the break-up of an older supercontinent, Columbia, created by global events of 2.0–1.8 Ga. Between 1300 and 900 Ma, Rodinia was established and the continental break-up occurred between 950 and 600 Ma. The first components of Gondwana begin to collide around the same time period, leading towards the creation of a global supercontinent.
2.
There are basically two geodynamic scenarios to explain the formation of the Gondwana Supercontinent in eastern and southern Africa during the Pan-African Orogenic Cycle:
1. i.
  collision and amalgamation of East Gondwana and West Gondwana (~640 to ~530 Ma) and closing of the Mozambique Ocean between 841 and 632 Ma as well as deformation during the East African Orogeny.
2. ii.
  collision and amalgamation of East Gondwana (ANS and older crystalline basement of the Dharwar Craton of southern India; Madagascar and the eastern granulites of Kenya and Tanzania), West Gondwana (Central Africa craton) and South Gondwana (Antarctica and the Kalahari craton).
3.
The Brasiliano Orogeny is the term used for the larger Pan-African/Brasiliano Orogeny that extended not only in South America but across most of Gondwana. The Damara–Zambezi–Lufilian Orogeny developed essentially by closure of linked, narrow ocean basins. The Zambezi Belt lies between the Congo Craton and the Kalahari Craton and branches off the MB in northern Zimbabwe and southern Zambia. The Damara Orogeny occurred late in the creation of Gondwana and the Damara Belt is exposed mainly in Namibia between the Kalahari and the Congo Cratons. The Lufilian Arc is about 800 km (500 mi) long extends across eastern Angola, the Katanga Province of the southern Congo and the northwest of Zambia. The Kaoko Belt branches north-west from the Damara Belt into Angola.
4.
The relics belonging to the East African Orogeny include high-grade metamorphic rocks within the Mozambique belt to greenschist facies rocks in NE Africa and the ANS. The late Mesoproterozoic–early Neoproterozoic events of the southern segment of the East African Orogeny include separation of the Congo and the contiguous East Sahara Cratons from Rodinia (~1200 Ma).
5.
The Kuunga Orogeny is an orogeny that occurred during the Ediacaran and Cambrian in SE Africa. It is slightly younger than the East African Orogeny and produced by collisions between India and Australia–East Antarctica and Azania and India.
6.
The ANS is southward narrowing belt straddling the suture separating East and West Gondwana at the northern end of the EAO. It extends over 3500 km north–south and more than 1500 east–west at its widest, underlying an area of ~2.7 × 10⁶ km² in the north half of the East African Orogeny. The ANS combines areas of western Arabia and northeastern Africa (Palestine, Jordan, Saudi Arabia, Egypt, Sudan, Ethiopia, Eritrea, Somalia and Yemen).
7.
Mid-Neoproterozoic ophiolites are common in the ANS, and range from 690 to 890 Ma and extend 3000 km N–S and > 1000 km E–W. A supra-subduction zone setting for the Egyptian Eastern Desert ophiolites is broadly popular but also a back-arc setting has also been inferred.
8.
Two stages for terrane accretion of the Arabian shield: (1) terrane accretion and (2) lateral transpressional accretion.
9.
The ANS took about 300 million years to form. The earliest rocks identified with the ANS crust’s formative process produced by the coalescence of island-arc and back-arc basins, and probably ocean plateaus. The oldest rocks of this series are around 870 Ma and are exposed in the eastern Sudan and South Arabia.
10.
The ENS covers ~100,000 km², crops out along the Red Sea Hills in the Eastern Desert and southern Sinai, as well as limited areas in the south Western Desert (Oweinat area). The basement complex of the ENS comprises four main tectonostratigraphic units from bottom upward as follows: (1) high-grade gneisses and migmatites, (2) arc-type volcanic/volcanosedimentary units, along with dismembered ophiolites, (3) the Ediacaran Hammamat and Dokhan supracrustal sequences, and (4) granitoids.
11.
The CED is dominated by ophiolitic rocks, arc volcanics and volcaniclastics as well as BIF and diamictite. The CED is mainly defined by the predominance of an NW-trending structural fabric that marks the NW–SE sinistral strike–slip shear zone related to Najd Fault System.
12.
The SED differs from the CED in several styles: (i) it is much less studied, (ii) it lacks interesting sediments such as BIF and diamictite, (iii) it has some volcanogenic massive sulphides, whereas the CED does not, (iv) lacks Hammamat molasse sediments and Dokhan Volcanics found in the NED and CED, and (v) has a similar abundance of serpentinite as the CED.
13.
The NED differs from the CED or the SED. It is characterized mainly by (i) abundance of Hammamat sediments and Dokhan Volcanics, (ii) absence of ophiolites and scarcity of gneisses, (iii) absence of BIF and diamictite, (iv) absence of Najd deformation, (v) abundance of bimodal dike swarms, (vi) abundance of epizonal A-type granites indicating strong extension, and (vii) abundance of Dokhan Volcanics.
14.
The basement rocks in Sinai are divided into four metamorphic complexes: the Feiran–Solaf, Sa’al–Zaghra, Kid and Taba Metamorphic Complexes. The metamorphic complexes contain orthogneisses and metasedimentary-metavolcanic sequences.

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Authors and Affiliations

Geology Department, Faculty of Science, Tanta University, Tanta, 31527, Egypt
Mohamed Abd El-Wahed
Department of Geology, Faulty of Science, Benha University, Benha, 13518, Egypt
Zakaria Hamimi

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Mohamed Abd El-Wahed
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Zakaria Hamimi
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Correspondence to Mohamed Abd El-Wahed .

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Department of Geology, Benha University, Benha, Egypt
Zakaria Hamimi
Department of Earth Sciences, Kanazawa University, Kanazawa, Japan
Shoji Arai
Geology Department, UAE University, Al Ain, United Arab Emirates
Abdel-Rahman Fowler
Geology Department, Port Said University, Port Said, Egypt
Mohammed Z. El-Bialy

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Abd El-Wahed, M., Hamimi, Z. (2021). The Egyptian Nubian Shield Within the Frame of the Arabian–Nubian Shield. In: Hamimi, Z., Arai, S., Fowler, AR., El-Bialy, M.Z. (eds) The Geology of the Egyptian Nubian Shield. Regional Geology Reviews. Springer, Cham. https://doi.org/10.1007/978-3-030-49771-2_2

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DOI: https://doi.org/10.1007/978-3-030-49771-2_2
Published: 30 September 2020
Publisher Name: Springer, Cham
Print ISBN: 978-3-030-49770-5
Online ISBN: 978-3-030-49771-2
eBook Packages: Earth and Environmental ScienceEarth and Environmental Science (R0)

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