Petrogenesis and Tectonic Evolution of Kab Amiri Ophiolites and Island-Arc Assemblages, Central Eastern Desert, Egypt: Petrological and Geochemical Constraints

Abstract

The Kab Amiri area in the Central Eastern Desert (CED) of Egypt comprises ophiolitic rocks, island arc metasediments, and granitic rocks. This study deals with the petrography and geochemistry of the ophiolitic and island arc rocks to understand their petrogenesis and geodynamic evolution of the CED ophiolitic belt. The Kab Amiri ophiolite is dismembered, comprising serpentinites and metabasalt (spilite). Serpentinites have low contents of Al₂O₃ (1.14 wt%) and CaO (0.65 wt%), suggesting they are depleted peridotite protoliths (e.g., depleted harzburgites to dunites). The high L.O.I. value (13.7 wt%) of serpentinite rocks indicates intense hydration and serpentinization during alteration processes. Petrographic and geochemical studies suggest that serpentinites were likely formed after depleted peridotites in a supra-subduction zone (SSZ) setting (e.g., a fore-arc setting). Spilitic basalt shows a tholeiitic affinity of the depleted mantle source. The arc-related metasediments are represented mainly by schists and slate. Many samples of metasediments are relatively low in alumina (Al₂O₃ < 15%), suggesting a low clay content and formation in an island arc setting. In contrast, protoliths of island-arc metabasalts and meta-andesites crystallized from calc-alkaline magmas in the immature oceanic arcs.

1. Introduction

The Eastern Desert of Egypt is divided into the northern (NED), central (CED), and southern (SED) sectors by two arbitrary lines that were drawn from Qena to Safaga and from Idfu to Marsa Alam (Figure 1a). The CED supracrustal sequences are an oceanic assemblage consisting of various Neoproterozoic ophiolitic rocks, primarily with greenschist facies, arc volcanic, as well as volcaniclastic rocks, banded iron formations (BIF), and diamictite. Ophiolitic ultramafic and island arc of the Neoproterozoic age are widespread in the CED and SED sectors of the Eastern Desert of Egypt [1,2,3]. Throughout the CED, there are many Neoproterozoic ophiolites. Figure 1a shows the distribution of various lithologies [3,4,5]. Ophiolite rocks of the Neoproterozoic age 800 and 730 Ma [6,7] are considered the most significant and widespread units with granitoid rocks in the Egyptian basement rocks of the Eastern Desert (ED) of Egypt. They are vestiges of oceanic lithosphere with intra-oceanic island arcs that are connected to the opening and closing of the Neoproterozoic Mozambique oceanic basin, and they constitute the northern extent of the East African Orogen 550–850 Ma [8,9]. Neoproterozoic ophiolites of the CED of Egypt are divided into the following categories: (i) complete MORB ophiolites; (ii) dismembered ophiolites; and (iii) arc-associated ophiolites [5].

Metasediments are abundant and extensively distributed in the Egyptian Eastern Desert south of the 26° N latitude line, both in the Central and Southern Eastern Desert and in southeastern Sinai. These metasediments exhibit substantial thicknesses, with values surpassing several kilometres and occupying substantial land areas. The basement complex in the Eastern Desert features distinctive ensimatic signatures, primarily composed of an island arc complex encompassing multiple banded iron formations, which are overlain by fragmented ophiolitic sequences and syn-tectonic gray granitoid [10,11,12]. The study area is primarily underlain by ophiolitic rocks (serpentinite and metabasalt), arc-related metasedimentary rocks, metavolcanic rocks (i.e., metabasalt and andesite), and different granitoid. The ophiolitic assemblages in the region are comprised of fragmented components such as serpentinite, metabasalt (spilite), and heterogeneous mélange rocks. Serpentinites, which are believed to have originated from fore-arc seafloor spreading, are the most widely occurring ophiolitic rock type in the area [2,13].

Additionally, younger alkali granites and older calc-alkaline granitoid have intruded into the island arc and ophiolitic sequences. Numerous faults trend in several directions, including NE–SW, NNW–SSE, NW–SE, WNW–ESE, and ENE–WSW in the region. While the NNW–SSW and NW–SE faults are frequently connected to significant left lateral displacement, the ENE–WSW and WNW–ESE faults allow minor right-lateral displacement [14]. In the investigated region, there are two primary generations of mineral foliation: older E–W foliation and younger, less prevalent N–S crenulation foliation, which is best developed in the metasediments and mélange matrix rocks. Several workers have studied the Kab Amiri area, such as [15,16,17,18,19,20,21]. Their research focused on the field and geophysical studies of stream sediments to ascertain the mineralization of radioactive elements. However, this research paper focuses on the basement rocks rather than stream sediments.

This study endeavours to clarify the geological and geochemical aspects of the study area by conducting a comprehensive examination of the petrology, alteration imprints and geochemistry of ophiolitic and island arc rock assemblages in the Kab Amiri area. The study aims to comprehend the tectonic environments and gain insight into these rocks’ petrogenesis.

Figure 1. (a) Simplified geologic map showing the classification of the eastern desert, Egypt, and location map of the Kab Amiri area (after compilation by Fowler cited in [22]. Abbreviation; Ad-Atmur-Delgo, CED-Central Eastern Desert, ED-eastern desert, Ga-Gabgaba, Gb-Gebeit, HHijaz, M-Midyan, NED-north eastern desert, S-Sinai, SED-south eastern desert, AH-Allaqi-Heiani, Ha-Hamisana, K-Keraf, OS-Onib-Sol Hamed, Y-Yanbu. (b) Geologic map of Kab Amiri area, Central Eastern Desert, Egypt (modified after [15].)

Figure 1. (a) Simplified geologic map showing the classification of the eastern desert, Egypt, and location map of the Kab Amiri area (after compilation by Fowler cited in [22]. Abbreviation; Ad-Atmur-Delgo, CED-Central Eastern Desert, ED-eastern desert, Ga-Gabgaba, Gb-Gebeit, HHijaz, M-Midyan, NED-north eastern desert, S-Sinai, SED-south eastern desert, AH-Allaqi-Heiani, Ha-Hamisana, K-Keraf, OS-Onib-Sol Hamed, Y-Yanbu. (b) Geologic map of Kab Amiri area, Central Eastern Desert, Egypt (modified after [15].)

2. Geologic Setting

Kab Amiri area is situated in the Central Eastern Desert of Egypt, south of the Qena-Safaga Road and mainly comprises Precambrian exposures. The main wadies draining in the Kab Amiri area include Wadis Kab Amiri, Abu Grahishi, El Bohlog, and Kab Um El Abas (Figure 1b). It is encircled by Latitudes 26°15′ and 26°25′ N and Longitudes 33°30′ and 33°40′ E. The lowest relief is south of Gabal Kab Amiri, while the highest relief (911 m above sea level) is at Gabal Kab Um El Abas, east of Gabal Kab Amiri. The volcano-sedimentary associations have been deformed and metamorphosed under greenschist to lower amphibolite facies conditions [23]. The ophiolitic rocks in the study area are represented by serpentinites and metabasalts (spilite). While island arc assemblages are comprised of metasediments and metavolcanic and followed by granitoid rocks and different post-orogenic dykes (pegmatite, basaltic) (Figure 2b). Several significant faults that cut across the Kab Amiri stretch mostly in the NW direction (Figure 1b) consist of the Najid fault trend, with a smaller number of faults trending in the NE and E–W directions [17,23].

Dismembered ophiolitic rocks are concentrated in the northern and central parts of the study area. It is noticeable in this region that the serpentinites turned into talc carbonate rocks (Figure 2a). Serpentinites are grey to greyish on the weathered surface, while the fresh surfaces have dark green. They have moderate to high relief. They vary in size and shape and always have tectonic contact with country rocks without thermal effects. The exposed serpentinites shape elongated bodies generally extending in the E–W direction southwest and southeast of Gabal Kab Amiri granites. Figure 2b shows a structural contact between serpentinites and metasediments. They are embedded in the metasediments mainly along axial planes of large anticlinal folds. The talc–carbonate rocks are found along faults and shear zones within the mountain range. Asbestos and calcite occur as veins about 5 cm thick along the fractures and shear zones, suggesting mixed H₂O-CO₂ introduction during deformation.

The Kab Amiri ophiolitic metabasalts are related to the older metavolcanic, which occupy substantial exposures in the northwestern sector of the study area, as depicted in Figure 1b. These metabasalts are elliptical (elongated) and irregularly curved pillow lobes, ranging in size from 40 to 70 cm in length, with a chilled margin visible along the peripheral boundary, as shown in (Figure 2c,d). The metabasalts are surrounded by a rim rich in epidote, chlorite, and other minerals, indicative of alteration processes on the seafloor and through subsequent hydrothermal activity, as depicted in Figure 2c. The metabasalts exhibit moderate topographic relief and are characterized by a dark grey to greenish-grey colouration, as depicted in (Figure 2d).

The metasediments and metavolcanic comprising the island arc assemblage constitute the majority of the rocks in the study area. The metasediments are primarily located in the southern portion of the mapped site, forming a curvilinear belt that encircles the Kab Amiri granite from the south, as depicted in Figure 1b. These rocks are characterized by medium-grained immature sediments, sometimes containing banded iron formations, and are primarily foliated, folded, and exhibit a greyish-green colouration with moderate to high relief. The lowest relief is observed in the southern region of Gabal Kab Amiri, while the highest relief, reaching 911 m above sea level, is present at Gabal Kab Um El Abas, located to the east of Gabal Kab Amiri. The metasediments are related to the Abu Fannani schist around the Meatiq dome [24]. The contacts between the metasediments and metavolcanic are tectonic in nature, particularly along Wadi El Saqia. The interaction between the dismembering ophiolite and Abu Fannani schist is tectonic, with the ophiolite rocks thrust over the Abu Fannani schist, which has been described as a tectonic mélange by [25]. These rocks exhibit an advanced schistosity and are bisected by quartz veins and veinlets, as depicted in Figure 2e.

The metavolcanic rocks are represented by substantial exposures located in the southern and northeastern parts of the mapped area and are characterized by tectonic contact with the metasediments. These rocks are related to the young metavolcanic (YMV), visible along Wadi El Saqia and extending east, as depicted in Figure 1b. These rocks form moderate to high relief ridges with dark grey to greenish-grey and pink colour. They are highly fractured, altered, sheared, and traversed by some mafic and felsic dykes (Figure 2f). The metavolcanic and metasediments are structural, while their contacts with Kab Amiri granites are intrusive.

3. Analytical Methods

In the study area, 12 samples were procured from ophiolitic rocks, 4 samples of serpentinites, 4 samples of talc-carbonates and 4 samples of metabasalts, and 11 samples from the island arc assemblage, including 7 samples of metasediments and 4 samples of metavolcanic. These samples were subjected to microscopic analysis using 15 thin sections prepared at the Laboratories of the Nuclear Materials Authority (NMA). A Nikon polarizing microscope (Nikon, Tokyo, Japan) equipped with an automatic photomicrographic and mechanical stage attachment was used to identify the minerals and textures.

Chemical analyses of whole rock samples were carried out at ACME Analytical Laboratories of Vancouver, Canada, for major oxides, trace, and rare earth elements. Total abundances of the major oxides and several minor elements are based on a 0.2 g sample analyzed by ICP-Optical emission spectrometry ICP-OES following a lithium metaborate/tetraborate fusion and dilute nitric digestion. Elements are expressed as common oxides for each element (i.e., Al₂O₃, CaO, Fe₂O₃, K₂O, MgO, MnO, Na₂O, P₂O₅, TiO₂). For determination of the loss on ignition (L.O.I.), 1 g of each sample was heated at 950 °C for 90 min, and then the weight loss during this process was calculated as L.O.I. The trace elements (Cr, Co, Sr, Zr, Nb, Mo, Cu, Pb, Zn, Ni, Bi and Sc) were determined by ICP-OES. In comparison, large ion lithophile elements (LILE) (Sr, Ba, Rb and Cs), high field strength elements (HFSE) (Zr, Nb, Y, Hf, Th, U, Ta) and rare earth elements (REE) (La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, and Lu) were analyzed by inductively coupled plasma mass spectrometry (ICP-MS) method at the same Laboratory (ACME). The concentration of major oxides and trace/rare earth elements are given in wt% (weight per cent) and parts per million (ppm), respectively. Detection limits for major oxides and trace elements vary from 0.001 wt% to 0.04 wt% and 0.01 to 0.5 ppm, respectively. Analytical precision calculated from replicate analyses was 0.5% for major oxides and varied from 2% to 20% for trace elements. The geochemistry (Minpet) program was used to calculate the current data in order to determine the classification, nomenclature, magma type, and tectonic setting of the granitic rocks in the study area. Furthermore, utilizing the major oxides and trace elements, some binary and ternary diagrams have been created for the same purpose.

4. Results

4.1. Petrography

Serpentinites are composed primarily of serpentine minerals (over 90% by volume), with minor amounts of other minerals such as olivine, chromite spinel, carbonates, and opaque minerals. The serpentine minerals antigorite, chrysotile, and lizardite are characterized by colourless, weakly birefringent, and parallel extinct crystals. Antigorite often occurs as elongated blade-like crystals, sometimes arranged in a parallel fashion, creating a pseudomorph after orthopyroxene minerals, known as the bastite texture. When the geometric configuration of the original mafic minerals is preserved, a pseudomorphic texture is created. In contrast, lizardite is a fine-grained matrix filling the pseudomorphs of olivine and is characterized by a mesh texture after olivine. Olivine is partially serpentinized and eventually altered to talc and carbonates. Chrysotile is a serpentine mineral present as randomly oriented fibrous crystals and associated with antigorite. Chromite is the primary opaque Mineral, forming subhedral fractured crystals filled with carbonates (Figure 3c).

Talc–carbonate rocks are massive and schistose a fine-grained texture. Their hues range from pale green to greenish-grey-coloured rocks and are often spotted with a brownish colour. These rocks comprise talc and carbonate with small quantities of opaque. Talc presents as fine-grained fibrous aggregates of parallel arrangement characterized by asbestos structure replacing the tremolite and actinolite (Figure 3d). It is colourless and displays strong birefringence with low relief, parallel extinction, and high interference colours. Talc also presents as a few crystals associated with the serpentine, while antigorite occasionally presents as blade-like crystals in the talc rocks (Figure 3e). The opaque are present as fine anhedral crystals of iron oxides disseminated throughout the rock (Figure 3f).

The ophiolitic metabasalts (spilite) are dark green consisting of plagioclase, chlorite, carbonate, and epidote. They are cryptocrystalline with an amygdaloidal texture that is oriented parallel to the schistosity planes. Plagioclase microphenocrysts are predominant in the rocks and occupy about 50%–70% of the studied rock volume. The crystals are intensively altered and albuminized, where sodium ions replace the calcium ions to form albite and liberate Ca form epidote and carbonate as secondary minerals (Figure 3g). Carbonates are common in this rock and constitute about 20 modal vol.% of the rock; they occur in two forms: as an aggregate of calcite associating the main constituents resulting from the process of albitization of the plagioclase (Figure 3h) and as amygdale fillings during seawater infiltration usually associated with the secondary quartz (Figure 3i). The rock primarily comprises chlorite, which makes up about 10% of the rock and is associated with plagioclase and mafic minerals.

The metasediments of Kab Amiri can be categorized according to the grain size into two types, schist and slate. These types are characterized by schistose texture, where the constituting minerals are elongated and foliated by stress.

Quartz–actinolite–chlorite schists are composed of amphibole, biotite, plagioclase, and quartz. Amphibole minerals are represented by hornblende, tremolite and actinolite (Figure 4a). Hornblende is found as hypidioblastic prisms commonly foliated and dark brown intensively altered to the tremolite-actinolite association and finally to chlorite. Actinolite is characterized by a brownish-green colour with pleochroism from pale green to greenish black. It is mostly associated with tremolite and is considered an alteration product of amphibole minerals. Tremolite, the most common amphibole, exists as fibrous crystals characterized by asbestos structure and exhibits third-order interference colours (Figure 4b). Biotite is less common and occurs as hypidioblastic pale brown flakes up to 0.4 mm long (Figure 4c). Sometimes, biotite occurs as porphyritic flakes surrounded by finer crystals of quartz, epidote, and carbonate (Figure 4c). Chlorite is more dominant and associated with mica minerals. It occurs as greenish folia with polysynthetic structure (Figure 4d) and as pseudomorphs after the earlier mafic minerals. Plagioclase is moderately saussuritized and altered to epidote and saussurite minerals, where its longer dimension is parallel to the schistosity (Figure 4c). Epidote is the most common secondary Mineral; it is found as oval crystals associating zoisite and chlorite as alteration products (Figure 4c). Quartz is found as elongated crystals with rounded to sub-rounded shapes. It also occurs as porphyritic elongate crystals with chlorite following the schistose texture (Figure 4d). The rock encloses a few crystals of primary calcite that are characterized by twinkling.

Slate is a fine-grained metamorphic rock produced by low-grade regional metamorphism for shale or mudstone, varying in colour from buff to brown and black. The main foliation is defined by aligned mafic minerals, such as chlorite showing well-developed schistosity. The slate is composed mainly of quartz, sericite, and chlorite. It appears banded due to the elongation and orientation of its constituents (Figure 4e). Quartz is the main constituent, comprising about 60 vol% of rocks and occurs as rounded to sub-rounded porphyroblastic crystals with diameters of about 0.6 mm (Figure 4e) or as minute crystals prevailing the groundmass. Chlorite exists as xenomorphic secondary flakes, associated with quartz, epidote and carbonates showing schistosity (Figure 4e). Chloritization is dominant, so the rock’s texture yields bands of green colour. Carbonates occur as cross-cutting micro veinlets of calcite (Figure 4f). Opaque minerals are represented by fine deep black grains disseminated through the rock. Sericite occurs as minute shreds and acicular crystals scattered throughout the groundmass (Figure 4e).

Petrographic investigation revealed that the metavolcanics are basic (metabasalts) and intermediate (metapyroclastic).

Metabasalts are fine-grained rocks with dark grey colour characterized by porphyritic texture. They are composed mainly of plagioclase and mafic minerals. Plagioclase occurs as hypidioblastic to xenoblastic crystals up to 0.6 mm in length and 0.3 mm in width (Figure 5a); it is partially altered to saussurite (Figure 5a) and epidote. Plagioclase also occurs as porphyritic crystals embedded in a fine-grained groundmass composed of plagioclase microlites and chlorite, exhibiting phenocryst texture. Mafic minerals are represented by actinolite, chlorite, and tremolite. The alteration minerals are also present as kinked xenoblastic crystals (up to 0.4 mm long) of chlorite tremolite (Figure 5b) and (Figure 5c) associated with epidote, iron oxides and relict biotite and hornblende. The mafic minerals are foliated and associated with carbonates as alteration products that are partially dissolved, replaced by secondary polycrystalline quartz (Figure 5a), and associated with epidote (Figure 5c).

The metapyroclastic rock is an andesitic and composed primarily of porphyritic clasts of plagioclase, quartz, carbonate, and perthite. The clasts are embedded in a fine-grained groundmass of chlorite and epidote. Plagioclase is the dominant Mineral, making up about 55% of the rock’s volume, and is present as porphyroblastic clasts and fine lathes in the groundmass. The plagioclase crystals have zoning and lamellar twinning and have been partially saussuritized (Figure 5d). The potash feldspar is present in the metapyroclastic rock as porphyritic clastic perthite with a string-like shape and measures 1.3 mm in length and 1.1 mm in width (Figure 5d), constituting about 20% of the rock. Few crystals of quartz mantled by plagioclase are recorded as porphyritic (Figure 5e). Quartz is also present as rounded to subrounded porphyroblastic crystals and fine crystals in the groundmass, accounting for about 15% modal vol.%. The porphyroblastic crystals are characterized by boundaries assimilated by the thermal effect of the groundmass (Figure 5e). Biotite and muscovite are found as minute flakes in the groundmass and altered to chlorite (Figure 5f), while epidote accounts for about 10% of the rock. The plagioclase crystals are present as porphyroblastic clasts and fine lathes in the groundmass, making up 55% of the rock and are characterized by zoning and lamellar twinning with partial saussuritization.

4.2. Geochemical Characteristics of the Ophiolitic Rocks

4.2.1. Chemical Compositions and Behaviour of Elements in the Ophiolitic Rocks

The dismembered ophiolitic rocks comprise serpentinites, talc–carbonate and metabasalts (spilite). The major oxides, trace and REEs elements of dismembered ophiolitic samples are recorded in

Table 1. The chemical composition of Kab Amiri investigated serpentinites, talc-carbonates and metabasalts.

Rock Type			Serpentinites                                             Talcite								Metabasalts
S. No.	9 C	20 D	10 C	16 E	Av.	19 D	17 E	9 E	23 D	Av.	6 B	11 B	17 B	6 E	Av.
	Major oxides (wt%)
SiO2	37.83	35.54	39.9	46.31	39.89	41.76	43.6	44.17	45.50	43.78	49.54	50.92	45.47	47.35	48.3
TiO2	0.01	0.01	0.02	0.07	0.03	0.16	0.09	0.35	0.32	0.23	1.15	1.47	0.98	1.18	1.2
Al2O3	0.62	0.45	0.50	3.00	1.14	3.98	5.69	4.32	4.50	4.62	14.70	15.25	13.19	13.49	14.16
FeOt	12.14	9.34	10.15	12.36	10.99	16.89	14.7	18.66	15.49	16.44	18.01	17.29	19.46	17.89	18.16
MnO	0.09	0.07	0.06	0.11	0.08	0.19	0.13	0.13	0.29	0.19	0.49	0.14	0.21	0.18	0.26
MgO	37.33	33.11	36.47	25.20	33.03	30.90	26.2	27.51	26.55	27.79	3.70	2.77	4.81	6.88	4.54
CaO	0.24	0.20	0.19	1.98	0.65	3.52	2.70	2.50	2.32	2.76	5.86	6.71	11.08	8.04	7.92
Na2O	0.02	0.01	0.01	0.02	0.02	0.41	0.09	0.18	0.05	0.18	2.70	3.98	2.12	2.70	2.88
K2O	0.01	0.01	0.01	0.01	0.01	0.24	0.10	0.22	0.16	0.18	0.04	0.40	0.12	0.18	0.19
P2O5	0.00	0.00	0.01	0.01	0.01	0.09	0.01	0.08	0.15	0.08	0.11	0.05	0.07	0.11	0.09
L.O.I.	11.11	20.76	12.34	10.43	13.66	1.53	6.19	1.48	4.57	3.44	3.45	1.00	2.39	1.57	2.10
Total	99.4%	99.5%	99.6%	99.5%	99.5%	99.7%	99.5%	99.6%	99.9%	99.7%	99.75%	99.98%	99.9%	99.57%	99.8%
Mg	0.75	0.80	0.78	0.67	0.75	0.65	0.64	0.60	0.63	0.63	-	-	-	-
Ol	71.7	69.2	70.6	69.1	-	65.6	70.1	72.5	69.7	-
Opx	24.8	28.4	27.8	26.9	-	28.9	25.9	21.3	25.5	-
Cpx	1.6	2.1	1.4	3.4	-	4.9	2.9	5.2	4.2	-
	Trace elements (ppm)
Mo	0.18	0.05	0.12	0.05	0.1	0.05	0.08	0.34	0.07	0.14	0.12	1.12	0.19	0.09	0.38
Cu	0.4	4.9	16.8	25.2	11.83	51.6	20.5	83.6	95.2	62.73	1.8	20.4	93.2	112.0	56.85
Pb	10.89	0.49	6.54	0.57	4.62	1.14	1.20	3.57	0.75	1.67	4.14	2.51	0.98	1.13	2.19
Zn	62.0	31.1	70.3	58.4	55.45	89.4	63.4	51.6	106.5	77.7	137.2	60.8	83.9	80.5	90.6
Ag	183	160	198	178	179	47	170	67.0	91	93.8	29.0	32.0	23.0	56.0	35
Ni	2130	2072	1948	1204	1838	240	1350	279.7	345	553	7.3	6.8	18.5	72.1	26.18
Co	101.4	93.2	110.2	84.4	97.3	39.8	90.5	33.7	62.5	56.6	29.7	20.7	40.6	46.9	34.48
As	0.4	1.9	2.3	3.7	2.1	2.4	2.9	10.2	2.4	4.5	1.4	1.1	1.5	1.5	1.38
U	0.1	0.1	0.2	0.1	0.1	0.1	0.1	0.2	0.1	0.1	0.1	0.1	0.1	0.2	0.13
Th	0.1	0.1	0.1	0.1	0.1	0.2	0.1	0.7	0.2	0.3	0.3	0.3	0.1	0.4	0.28
Sr	5.0	2.0	15.3	53.0	18.8	97	77.3	22	6.0	50.6	125.0	123.0	138	174	140
Cd	0.06	0.10	0.09	0.10	0.09	0.10	0.10	0.10	0.02	0.08	0.12	0.08	0.10	0.04	0.09
Sb	0.08	0.58	0.40	0.11	0.29	0.61	0.10	0.40	0.47	0.39	0.20	0.04	0.19	0.15	0.15
Bi	0.04	0.06	0.05	0.04	0.05	0.04	0.04	0.04	0.04	0.04	0.07	0.04	0.04	0.04	0.05
V	32.0	24.0	50.4	62.0	42.1	325	102	133	541	275	312	164	379	288	285.8
Cr	1459	1487	1873	1989	1702	485	1830	428	542	821	10.0	8.0	20.0	97.0	33.75
Ba	1.0	1.0	2.5	3.0	1.9	25.0	2.0	4.3	32.5	15.9	26.0	21.0	20.0	22.0	22.25
W	1.7	0.1	0.9	0.1	0.7	0.1	0.1	0.1	0.4	0.2	0.3	0.1	0.1	0.1	0.15
Zr	1.0	0.4	1.6	5.0	2	26.9	4.2	16.5	5.2	13.2	26.4	15.3	24.3	27.7	23.42
Sn	0.4	0.1	0.6	1.0	0.5	0.8	0.8	0.4	0.7	0.7	0.7	3.0	0.5	0.7	1.23
Be	1.0	1.0	1.2	1.0	1.05	1.0	1.0	1.0	1.0	1.0	1.0	1.0	1.0	1.0	1.0
Sc	3.4	4.3	4.6	12.5	6.2	36.5	17.5	26.3	57.4	34.4	32.5	23.2	42.0	43.2	35.23
S	0.10	0.10	0.09	0.04	0.08	0.10	0.10	0.04	0.04	0.07	0.10	0.10	0.10	0.10	0.1
Y	0.3	0.4	0.9	2.2	0.9	23.5	1.6	9.0	4.0	9.5	31.1	20.0	22.5	21.9	23.88
Hf	0.05	0.02	0.02	0.16	0.06	1.12	0.18	0.65	0.14	0.52	0.5	0.58	0.99	1.04	0.78
Li	1.4	1.3	1.9	4.7	2.3	10.4	1.5	6.1	18.9	9.2	25.1	13.3	7.8	20.8	16.75
Rb	0.5	0.1	0.2	0.1	0.2	1.8	0.1	3.9	4.3	2.5	0.5	3.8	1.6	2.7	2.15
Ta	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.1	0.2	0.13
Nb	0.08	0.04	0.10	0.21	0.11	0.79	0.17	1.87	1.19	1.01	1.27	1.13	0.72	1.04	1.04
Cs	0.2	0.1	0.2	0.1	0.15	0.1	0.1	0.1	0.3	0.2	0.1	0.2	0.1	0.1	0.13
Ga	1.80	0.85	2.47	3.06	2.05	13.60	2.80	12.87	20.22	12.37	20.15	13.39	16.92	13.57	16.1
In	0.02	0.01	0.02	0.02	0.02	0.07	0.02	0.04	0.13	0.07	0.10	0.10	0.08	0.07	0.09
Re	0.002	0.002	0.002	0.002	0.002	0.002	0.002	0.002	0.002	0.002	0.002	0.002	0.003	0.006	0.003
Se	0.4	0.4	0.05	0.7	0.39	0.3	0.5	0.3	0.3	0.4	0.3	0.3	0.3	0.3	0.3
Te	6.50	5.75	7.11	4.71	6.02	0.65	4.0	1.81	0.74	1.80	0.34	0.30	0.81	1.59	0.76
Tl	0.05	0.05	0.05	0.05	0.05	0.05	0.05	0.5	0.22	0.21	0.05	0.06	0.05	0.05	0.05
	Rare earth elements REEs (ppm)
La	0.1	0.1	0.3	0.9	0.35	2.4	0.8	5.2	4.1	3.1	3.1	2.6	1.6	4.6	2.98
Ce	0.25	0.25	0.40	1.57	0.62	6.88	1.25	10.8	12.1	7.8	8.55	7.23	4.76	11.70	8.06
Pr	0.1	0.1	0.2	0.2	0.2	1.3	0.3	1.7	2.0	1.3	1.5	1.1	1.0	1.8	1.35
Nd	0.2	0.2	0.6	0.8	0.5	6.5	1.0	5.5	9.9	5.7	7.8	5.3	5.1	8.6	6.7
Sm	0.1	0.1	0.1	0.3	0.2	2.5	0.3	1.4	2.8	1.8	2.9	1.6	2.0	2.9	2.35
Eu	0.1	0.1	0.2	0.1	0.1	0.7	0.1	0.4	0.6	0.5	1.3	0.6	0.9	1.1	0.98
Gd	0.1	0.1	0.1	0.1	0.1	3.0	0.2	1.7	2.1	1.8	4.0	2.4	3.0	3.6	3.25
Tb	0.1	0.2	0.2	0.1	0.2	0.6	0.1	0.4	0.2	0.3	0.8	0.5	0.5	0.6	0.6
Dy	0.1	0.1	0.2	0.3	0.2	4.0	0.2	1.7	1.0	1.7	4.8	2.9	4.0	4.3	4.0
Ho	0.1	0.1	0.1	0.1	0.1	1.0	0.1	0.4	0.1	0.4	1.2	0.8	1.0	0.9	0.98
Er	0.1	0.1	0.1	0.3	0.2	2.5	0.4	1.0	0.4	1.1	3.3	2.2	2.4	2.3	2.55
Tm	0.1	0.1	0.2	0.1	0.1	0.4	0.1	0.1	0.1	0.2	0.5	0.4	0.4	0.4	0.43
Yb	0.1	0.2	0.1	0.2	0.2	2.4	0.1	0.9	0.4	0.9	3.0	2.3	2.1	2.1	2.38
Lu	0.1	0.1	0.2	0.1	0.1	0.4	0.1	0.1	0.1	0.2	0.5	0.4	0.4	0.3	0.4
Nb/La	0.80	0.4	0.33	0.23	0.44	0.32	0.21	0.36	0.29	0.29	0.41	0.43	0.45	0.23	0.38
Nb/Ce	0.32	0.16	0.25	0.13	0.21	0.11	0.14	0.17	0.10	0.13	0.15	0.16	0.15	0.09	0.14
Ce/Pb	0.02	0.51	0.02	2.75	0.83	6.04	1.04	3.03	16.12	6.56	2.07	2.88	4.86	10.35	5.04
Nb/U	0.8	0.4	0.5	2.1	0.9	7.9	1.7	9.35	11.9	7.7	12.7	11.3	7.2	5.2	9.1
La/Sm	1	1	3	3	2	0.96	2.67	3.71	1.46	2.2	1.07	1.63	0.8	1.59	1.27
Lu/Yb	1	0.5	2	0.5	1	0.17	1	0.11	0.25	0.38	0.17	0.17	0.19	0.14	0.17
ΣREE	1.65	1.65	3.0	5.17	2.87	34.58	5.05	31.32	35.89	26.71	43.25	30.33	19.16	45.2	34.49
LREE	0.85	0.85	1.80	3.87	1.84	20.3	3.75	25.02	31.49	20.14	25.15	18.43	15.36	30.7	22.41
HREE	0.80	1.0	1.20	1.3	1.1	14.3	1.3	6.3	4.4	6.6	18.1	11.9	13.8	14.5	14.58
LREE/HREE	1.1	0.85	1.5	3.0	1.6	1.4	2.9	4.0	7.2	3.9	1.4	1.5	1.1	2.1	1.53
Eu/Eu*	3.0	3.0	2.5	1.2	2.4	0.8	0.8	3.0	0.8	1.4	1.2	0.9	1.1	1.0	1.05
La/YbN	0.7	0.3	0.6	1.4	0.8	0.7	1.2	0.8	0.7	0.9	0.7	0.8	0.5	1.5	0.88
Tb/YbN	4.3	4.3	4.1	3.0	3.9	1.1	5.4	3.9	6.9	4.3	1.1	0.9	1.0	1.2	1.05
Gd/YbN	0.8	0.4	0.9	2.1	1.1	1.0	4.3	1.9	2.1	2.3	1.1	0.8	1.2	1.4	1.13
La/SmN	0.6	1.2	1.1	1.1	1	0.6	1.1	3.5	4.2	2.3	0.7	1.0	0.5	1.0	0.8
Ce/Ce*	0.949	0.949	0.915	0.867	0.920	0.983	0.693	0.976	1.051	0.926	0.969	1.048	0.966	0.987	0.99

FeO^t is total iron oxides (Fe₂O₃ + FeO).

. Serpentinite samples exhibit high enrichment in MgO, L.O.I., Ni, Co, and Cr contents relative to metabasalt (spilite). On the other hand, metabasalt (spilite) samples are enriched in SiO₂, Al₂O₃, FeOt, CaO, Na₂O, K₂O, Zn, V, Ba, Zr, Y, HREE, and LREE compared with serpentinites and talc-carbonates (Table 1). The average SiO₂ content varies in the range from 35.54 to 46.31 wt% due to different serpentine phases from antigorite (high SiO₂ up to 46.31 wt%) to lizardite with low SiO₂ (around 35.54 to 37 wt%) and has a low amount of Al₂O₃ 0.45–3.0 wt%, and CaO 0.20 to 1.98 wt% and high total FeO^t 9.34–12.36 wt%, suggesting depleted harzburgite protoliths [1,26,27]. Furthermore, the majority of MnO, Na₂O, K₂O, TiO₂ and P₂O₅ contents in the serpentinites are lower than 1 wt% and show the depletion degree of the mantle of investigated samples [28]. Volatile components are reported as a loss on ignition (L.O.I.) and can be used to measure the degree of serpentinization. The average L.O.I. content varies from 10.43 to 20.76 wt%, with an average of 13.7 wt%. The high value of L.O.I. in serpentinite rocks reflects intense hydration and serpentinization during alteration processes and abundance in talc and carbonates. According to [29], L.O.I. values of more than six wt% are regarded as changed because either (1) these samples are partially dehydrated during their subduction and prograde metamorphism or (2) a higher amount of antigorite.

Trace element analyses show that the samples tend to be richer in Cr (1989 ppm), Ni (2130 ppm) and Co (33.7–101.4 ppm) than the other analyzed trace elements. According to [30], serpentinites have a higher concentration of lithium than saltwater, which is consistent with this property [31]. Bulk serpentinites have lithium contents ranging from 1.3 to 18.9 ppm. According to [32,33], the processes that took place before serpentinization may have contributed to the Li enrichment, and serpentinization itself ought to lower the bulk sample’s Li concentration. The high mobile arsenic, antimony, and lead in fluids made these elements represent potential tracers of the nature of fluids during serpentinization. The serpentinites show (0.4 < As < 10.2 ppm; 0.08 < Sb < 1 ppm)that they are moderate to highly enriched in these elements compared to the primitive and depleted mantle (0.1 < As < 10 ppm; 0.001 < Sb < 1 ppm) [34,35]. They showed that they are transferred by aqueous fluids from the slab to the mantle wedge and are incorporated into serpentinites under oxidized conditions at shallow depths (~25 km) [36]. The serpentinites show Pb concentrations from 0.49 to 10.89 ppm. Nb/La and Nb/Ce value ranges of the serpentines are (0.23–0.80) (0.13–0.32), with averages (of 0.44, 0.21), respectively. These values are lower than that of the primitive mantle (PM 1.02 and 0.40, respectively [33,37], 1.04 and 0.40, respectively, [38] and the average bulk crust (0.69 and 0.33, respectively), which could be evidence of possible crustal contamination. Mantle-derived magmas are characterized by high Ce/Pb (25 ± 5) and Nb/U (47 ± 10), low Lu/Yb (0.14–0.15), and relatively lower La/Sm (<4.5), [39]. Continental crust has relatively low Ce/Pb (<15) and Nb/U (~9.7), along with higher Lu/Yb (0.16–0.18) and La/Sm (>4.5). Thus, these ratios can be used to reveal crustal contamination. All samples of serpentinites samples have a La/Sm ratio of <4.5, indicating a mantle derived source [39]. In addition, the Nb/U ratio is lesser than 9.7, in addition, all serpentinite samples have Ce/Pb ratio of <15, which also indicates some crustal contamination.

4.2.2. Protolith of Serpentinites

According to the Ol-Opx-Cpx ternary diagram [40], the serpentinites and talc-carbonates are mainly formed after harzburgites, with minor dunites (Figure 6a) because of the serpentinites alteration and lack of primary mineral phases, low content of Al₂O₃ and CaO and higher MgO. The (AFM) (Na₂O + K₂O)-FeO^t-MgO diagram [41] can discriminate tholeiitic and calc-alkaline affinity. Plotting the investigated serpentinites and talc-carbonates on this diagram revealed that the samples plot within the calc-alkaline field (Figure 6b). The most significant variations among the major oxides are usually in the abundances of K₂O and SiO₂. According to the binary diagram of SiO₂ and K₂O of [42], the serpentinites and talc-carbonates plot in the low-K tholeiitic field (Figure 6c).

5. Discussion

5.1. Tectonic Implication of Serpentinites and Talc-Carbonates

The studied serpentinites and talc-carbonates are low in Ti and high in Cr# [19] (Figure 7a,b). The Cr-spinel in serpentinites lies in fore-arc peridotites [2,27], but Cr-spinel in dunites and some highly depleted harzburgites plots in the boninite field. In the Cr# [19] and Mg# diagram (Figure 7c), the serpentinites and talc-carbonates lie in fore-arc peridotites, showing a good negative Cr# − Mg# trend reflecting the partial melting trend from harzburgites to dunites. Based on the variation diagram of [43] of SiO₂/MgO versus Al₂O₃, the serpentinites plot in the peridotite field (Figure 7d).

In the MgO/SiO₂ and Al₂O₃/SiO₂ diagrams (Figure 8a), each sample examined plotted below the “terrestrial mantle array” trend and showed a trend of magmatic depletion or enrichment from a primitive mantle to a highly depleted harzburgitic composition [44,45,46,47]. This shift could be attributed to Mg loss due to seafloor weathering during serpentinization [45,48]. The Al₂O₃/SiO₂ and MgO/SiO₂ ratios are comparable to those of Arabian shield and fore-arc peridotites (Figure 8a; [22,27,49]). When plotting compiled serpentinites and talc-carbonates in an Al₂O₃ versus CaO (wt%) diagram (Figure 8b) of [27,49] all samples plotted in the fore-arc peridotite field. The serpentinites have low Al₂O₃ and CaO contents, like depleted fore-arc peridotite [2,27]. On the SiO₂/MgO versus Al₂O₃ diagram (Figure 8c), they are affiliated with ophiolitic peridotites, as are the other Eastern Desert ophiolitic ultramafics [43]. The investigated serpentinites differ significantly from those found elsewhere in Egypt, as seen in Figure 8c. The Al₂O₃ content seems to be relatively uninfluenced by serpentinization, demonstrating that the bulk-rock Al content typically reflects its original primary concentration [28,45]. Low Al₂O₃ abundances (average 1.14 wt%) define the investigated serpentinites, like peridotites from the fore-arc setting and Neoproterozoic serpentinized peridotites from the Eastern Desert, Egypt (Figure 8d). The chemistry of serpentinites (low in Al₂O₃, CaO and TiO₂, but high in MgO, Ni, Cr and Co) suggests that they were formed after depleting harzburgite protoliths in fore-arc settings.

Figure 7. (a) Cr# after [19] versus TiO₂ binary diagram [50] of the serpentinites and talc-carbonates. (b) TiO₂ versus Cr# diagram [50,51]. (c) Mg# versus Cr# binary diagram [52]. CED field for chromites in the Central Eastern Desert serpentinites is from [53]. (d) SiO₂/MgO versus Al₂O₃ binary diagram [43]. Symbols as in Figure 6.

Figure 7. (a) Cr# after [19] versus TiO₂ binary diagram [50] of the serpentinites and talc-carbonates. (b) TiO₂ versus Cr# diagram [50,51]. (c) Mg# versus Cr# binary diagram [52]. CED field for chromites in the Central Eastern Desert serpentinites is from [53]. (d) SiO₂/MgO versus Al₂O₃ binary diagram [43]. Symbols as in Figure 6.

Figure 8. (a) MgO/SiO₂ versus Al₂O₃/SiO₂ of the serpentinites and talc-carbonates. Depleted mantle, primitive mantle, and abyssal peridotite are from [43,54,55], respectively. The black line is the terrestrial array after [46]. The blue area highlights a global compositional range of abyssal peridotites [43]. (b) Al₂O₃ versus CaO binary diagram of [49]. (c) SiO₂/MgO ratios versus Al₂O₃ diagram. Fields of ophiolitic gabbros and peridotites as well as MORB are from [43]. Data from the Eastern Desert are shown for comparison [56,57]. (d) The contents of the bulk-rock Al₂O₃ (wt%) of Kab Amiri serpentinites were compared with those from different tectonic settings and the Pan-African serpentinites [58]. Symbols as in Figure 6.

Figure 8. (a) MgO/SiO₂ versus Al₂O₃/SiO₂ of the serpentinites and talc-carbonates. Depleted mantle, primitive mantle, and abyssal peridotite are from [43,54,55], respectively. The black line is the terrestrial array after [46]. The blue area highlights a global compositional range of abyssal peridotites [43]. (b) Al₂O₃ versus CaO binary diagram of [49]. (c) SiO₂/MgO ratios versus Al₂O₃ diagram. Fields of ophiolitic gabbros and peridotites as well as MORB are from [43]. Data from the Eastern Desert are shown for comparison [56,57]. (d) The contents of the bulk-rock Al₂O₃ (wt%) of Kab Amiri serpentinites were compared with those from different tectonic settings and the Pan-African serpentinites [58]. Symbols as in Figure 6.

5.2. Classifications and Magma Type of Basalts

The present samples show sub-alkali basalt affinity on the Zr/TiO₂ × 0.0001 vs. Nb/Y diagram [59] (Figure 9a). While on the variation diagram of (Na₂O + K₂O) vs. SiO₂ [60], all samples plot in the subalkaline field (Figure 9b). The Ti/Y versus Nb/Y ratio diagram discriminates between the tholeiitic and alkaline magmas [61]. The metabasalts samples lie in the tholeiitic field (Figure 9c). These features indicate that the parental magma of basalt was tholeiitic in composition.

5.3. Tectonic Significance and Petrogenesis of Basalts

The geochemical characteristics of the ophiolite suite of rocks, in general, show variations in their major and trace elemental concentration, which could be due to mantle heterogeneity without any influence of crustal contamination [45,62]. In the Th/Yb vs. Nb/Yb discrimination diagrams of [63], all the metabasalts rocks fall above the N-MORB array suggesting arc-related magmatism (Figure 9d). Figure 10a shows that there is considerable overlapping between N-MORBs and E-MORBs, and between E-MORBs, P-MORBs, and ABs. Th and Nb behave similarly during both partial melting and fractional crystallization processes (see trends in Figure 10b). In the Th-Nb diagram, basalts generated in oceanic subduction-unrelated settings, rifted margins, and OCTZ can be distinguished from subduction-related basalts with a misclassification rate of <1% (Figure 10b). Three different types of convergent plate settings can be discriminated on the Th_N vs. Nb_N diagram in Figure 10b. Island arcs with complex polygenetic crustal nature are primarily characterized by the occurrence of CABs, which are displaced to the highest Th- Nb values. Intra-oceanic arcs display a large variability in Th-Nb contents, which can be used for recognizing two sub-types of intra-oceanic arc basalts. The Th/Nb enrichment indicates subduction-mantle source interaction, whereas decreasing Th–Nb compositions concerning PM define an array of mantle depletion without contribution from subduction-derived components Figure 10a,b.

Chondrite-normalized rare Earth elemental (REE) patterns of the metabasalts are shown in (Figure 11). The studied rocks have rare earth element (REE) contents (ΣREE = 19.16–45.2 ppm). These metabasalts show moderate LREE/HREE fractionation (La/Yb)_N = 0.5 to 1.5 and (Gd/Yb)_N = 0.8 to 1.4 with Eu/Eu* = 0.9–1.2 and display flat pattern without an Eu anomaly, similar to typical N-MORB.

Figure 11. Chondrite-normalized REEs diagram of [64] for the studied metabasalts.

Figure 11. Chondrite-normalized REEs diagram of [64] for the studied metabasalts.

5.4. Geochemical Behavior of Elements in Metasediments

5.4.1. Elemental Distribution in Schists

The chemical analyses of seven representative samples of the island arc metasediments and four representative samples of the island arc metavolcanic of the Kab Amiri area are listed in

Table 2. Whole-rock major oxides, trace and REEs element compositions of metasediments and metavolcanic rocks from Kab Amiri.

Rock Type	Metasediments								Metavolcanic
S. No.	5 D	14 D	7 E	18 E	4 D	2 E	3 F	Av.	6 D	17 D	3 D	11 D	Av.
Major oxides (wt%)
SiO2	51.10	57.78	51.77	49.20	53.10	59.69	64.52	55.3	56.63	54.10	55.10	51.95	54.45
TiO2	0.39	0.34	0.30	1.42	0.99	1.23	0.62	0.75	0.88	0.73	1.31	0.92	0.96
Al2O3	12.57	13.42	15.44	14.99	15.68	14.85	16.36	14.76	12.50	12.61	13.0	13.81	12.98
FeOt	11.74	12.14	16.38	12.36	9.5	7.62	4.32	10.56	9.12	10.69	8.95	9.01	9.41
MnO	0.14	0.11	0.13	0.17	0.19	0.14	0.13	0.14	0.31	0.33	0.21	0.14	0.24
MgO	6.09	6.58	5.60	7.30	6.69	3.21	1.03	5.2	9.10	8.50	8.99	10.0	9.14
CaO	7.08	4.87	6.91	9.6	8.7	5.43	1.32	6.26	8.0	8.72	7.90	9.0	8.4
Na2O	2.83	2.43	1.93	2.1	3.2	4.89	6.38	3.38	2.20	2.90	2.85	3.4	2.83
K2O	0.28	0.18	0.76	0.19	0.14	1.14	3.89	0.9	0.92	1.0	1.15	0.91	0.98
P2O5	0.05	0.07	0.05	0.10	0.11	0.09	0.09	0.08	0.10	0.08	0.12	0.09	0.09
L.O.I.	7.53	2.02	0.54	2.45	1.46	1.31	1.14	2.32	0.14	0.15	0.27	0.48	0.39
Total	99.8	99.94	99.81	99.88	99.76	99.6	99.8	99.65	99.9	99.81	99.85	99.71	99.87
Trace elements (ppm)
Mo	0.16	0.8	0.17	0.52	0.47	0.36	0.27	0.39	0.16	0.19	0.12	0.17	0.16
Cu	54.9	18.0	77.9	89	97	18.5	23.5	54.11	98.5	80.9	37.5	30.2	61.78
Pb	1.51	1.83	2.16	2.9	3.1	4.1	2.8	2.63	3.90	4.50	4.17	3.12	3.92
Zn	57.1	56.4	122.5	83	96.3	98.7	78.9	84.7	77.1	100.7	99.3	80.6	89.43
Ag	45.0	45.0	30.0	30.0	50.3	35.6	40.5	39.48	29.0	50.7	41.2	33.5	38.6
Ni	19.9	12.4	60.2	94	65	18	23	41.79	80.2	47.5	12.6	50.3	47.65
Co	21.1	17.6	30.0	41	38	21	17.5	26.6	39.8	59.3	51.5	49.2	49.95
As	1.0	0.7	0.8	0.8	0. 7	0.9	1.1	0.76	3.0	2.9	3.2	2.1	2.8
U	0.2	0.1	0.1	0.3	0.4	0.2	0.4	0.24	0.2	0.1	0.2	0.1	0.15
Th	0.3	0.4	0.3	0.4	0.6	0.2	0.5	0.39	0.4	0.3	0.4	0.5	0.4
Sr	102.0	87.0	258	160.3	139	140	186	153.19	340.0	195.5	244.1	207.9	246.9
Cd	0.13	0.07	0.09	0.10	0.9	0.7	0.19	0.31	12.3	9.3	6.9	7.5	9.0
Sb	0.24	0.06	0.51	0.27	0.09	0.9	0.21	0.33	0.17	0.09	0.03	0.10	0.09
Bi	0.05	0.04	0.05	0.03	0.04	0.08	0.07	0.05	0.04	0.03	0.06	0.07	0.05
V	112	144	245	241	342	226.9	175.6	212.36	176	290	312	210	247
Cr	120	58.0	17.0	145	75.7	19.6	21.9	65.31	33.0	22.0	17.5	10.0	20.62
Ba	72.0	27.0	283	36.2	48.7	3.98	4.12	67.86	105.0	85.0	96.5	100.7	96.8
W	0.1	0.1	0.2	0.1	0.2	0.4	0.3	0.2	0.2	0.3	0.6	0.4	0.375
Zr	40.2	28.3	15.2	51.8	98.3	75.9	100.3	58.57	98.6	82.6	70.2	56.3	76.92
Sn	0.5	0.5	0.3	0.6	0.9	0.5	0.7	0.57	0.4	0.6	1.8	0.8	0.9
Be	1.0	1.0	1.0	1.0	1.0	1.0	1.0	1.0	1.0	0.9	1.0	1.0	0.98
Sc	22.3	21.8	29.3	25.1	30.2	28.7	19.8	25.31	36.5	29.5	41.2	30.3	34.38
S	0.10	0.10	0.10	0.10	0.10	0.10	0.10	0.1	0.10	0.10	0.10	0.10	0.1
Y	14.6	13.6	6.2	21.8	19.3	18.9	8.6	14.71	29.3	27.5	29.3	32.0	29.53
Hf	0.94	0.64	0.74	3.2	5.9	6.3	5.7	3.35	0.83	1.2	0.92	0.61	0.89
Li	10.7	27.6	27.6	12.0	18.9	32.5	19.8	21.3	9.7	18.2	15.3	7.5	12.67
Rb	6.0	13.8	13.6	3.6	4.2	30.8	15.7	12.53	17.8	20.3	18.6	20.2	19.22
Ta	0.1	0.1	0.1	2.0	1.8	3	1.7	1.26	0.2	0.1	0.4	0.2	0.23
Nb	0.77	0.67	0.69	2.9	2.1	7	4.3	2.63	3.7	2.9	3.1	2.5	3.05
Cs	0.1	0.8	10.9	0.7	0.7	0.9	0.8	0.7	0.2	0.1	0.3	0.1	0.18
Ga	10.38	11.69	12.69	16.3	12.9	30.4	27.3	17.38	22.7	18.9	19.2	15.5	19.08
In	0.04	0.03	0.03	0.03	0.03	0.03	0.03	0.03	0.04	0.06	0.09	0.10	0.07
Re	0.002	0.003	0.003	0.003	0.003	0.002	0.003	0.002	0.002	0.002	0.002	0.002	0.002
Se	0.3	0.3	0.3	0.4	0.3	0.5	0.3	0.34	0.33	0.29	0.30	0.21	0.28
Rare earth elements REEs (ppm)
La	2.6	2.4	2.0	4.3	4.9	5.1	4.9	3.74	2.9	3.8	7.2	4.5	4.6
Ce	6.83	5.95	4.38	5.2	6.1	5.6	4.3	5.48	8.5	12.1	14.7	10.3	11.4
Pr	1.1	3.8	1.2	2.1	1.1	2.0	1.7	1.86	1.7	1.5	1.3	1.0	1.38
Nd	5.2	4.9	2.6	3.7	4.2	2.9	3.2	3.81	9.7	8.2	8.1	7.5	8.38
Sm	1.7	1.5	0.7	1.3	1.2	0.9	1.5	1.26	5.8	4.7	1.9	3.5	3.98
Eu	0.5	0.5	0.2	0.4	0.3	0.2	0.7	0.4	1.2	1.1	0.8	0.9	1.0
Gd	2.2	2.2	0.9	1.9	2.6	3.0	1.9	2.1	6.3	8.9	5.2	7.0	6.85
Tb	0.4	2.9	0.6	0.4	3.5	0.9	1.0	1.39	0.5	0.7	0.9	0.8	0.73
Dy	2.5	2.1	1.0	1.3	2.5	2.2	2.5	2.01	4.0	4.1	3.9	4.7	4.18
Ho	0.6	0.5	0.2	0.3	0.7	0.2	0.3	0.4	0.8	0.9	1.2	1.3	1.05
Er	1.7	1.6	0.7	1.9	1.6	1.5	1.8	1.54	6.0	4.1	5.2	3.9	4.8
Tm	0.3	0.2	0.1	0.4	0.3	0.4	0.3	0.29	0.4	0.5	0.8	0.2	0.48
Yb	1.5	1.4	0.8	1.6	0.4	1.9	0.8	1.2	3.6	4.3	5.2	6.0	4.78
Lu	0.2	0.3	0.1	0.2	0.2	0.3	0.5	0.26	0.5	0.6	0.3	0.2	0.4
ΣREE	27.3	30.1	15.5	25	29.6	27.1	25.4	25.71	51.9	55.5	56.7	51.8	53.98
LREE	17.9	19	11	17	17.8	16.7	16.3	16.53	28.9	31.4	34.0	27.7	30.5
HREE	9.4	11.2	4.4	8	11.8	10.4	9.1	9.19	21.1	24.1	22.7	24.1	23
LREE/HREE	1.9	1.7	2.5	2.1	1.5	1.6	1.8	1.87	1.4	1.3	1.5	1.4	1.4
Eu/Eu*	0.8	0.8	0.8	0.8	0.5	0.3	1.3	0.76	0.6	0.5	0.7	0.5	0.58
La/YbN	1.2	1.2	1.7	1.8	8.3	1.8	4.1	2.87	0.5	0.6	0.9	0.5	0.63
Tb/YbN	1.1	8.9	3.2	1.1	37.4	2.0	5.3	8.43	0.6	0.6	0.7	0.6	0.63
Gd/YbN	1.2	1.3	0.9	1.0	5.3	1.3	1.9	1.84	1.4	1.7	0.8	0.9	1.2
La/SmN	1.0	1.0	1.8	2.1	2.6	3.6	2.1	2.03	0.3	0.5	2.4	0.8	1.0
Ce/Ce*	0.997	0.934	0.959	0.608	0.627	0.634	0.484	0.75	0.937	1.178	0.938	0.923	0.99

. The metasediments are characterized by high silica contents from 49.20 to 64.52 wt% and A1₂O₃ 12.57 to 16.36%, which reflects the dominance of aluminous clay minerals. The percentages of MgO, CaO, and Na₂O in metasediments range from 1.03 to 7.30, 1.32 to 9.6, and 1.93 to 6.38, respectively. The Ni, Cr, Sc, and V elements, regarded as compatible ferromagnesian trace elements, exhibit quite variable abundances in the metasediments. Ni is in the range from 18 to 94 ppm, while Cr content varies from 17 to 145 ppm. Alternatively, Sc and V abundance ranges are from 19.8 to 30.2 ppm and 112 to 342 ppm, respectively. The Large ion lithophile elements (LILEs) Cs, Rb, Ba, K and Sr concentrations are variable. Cs have lower content and vary from 0.1 to 0.9 ppm, Rb ranges from 3.6 to 30.8 ppm, and Ba from 3.98 to 2.83 ppm. Sr is depleted and ranges from 87 to 258 ppm. Depletion of Sr is connected to the low abundance of CaO, implying that these metasediments’ source material is plagioclase poor. Th, U, Zr and Y show concentrations ranging from 0.2 to 0.6 ppm, 0.1 to 0.4 ppm, 15.2 to 100.3 ppm, and 6.2 to 21.8 ppm, respectively. In addition, normal to low concentration of Zr, which is concentrated in zircon and increases with maturity, infers immaturity to semi-maturity for these metasediments (Table 2). The related pairs Nb and Ta show exceptional coherence and are depleted in all samples, and they range from 0.67 to 7.0 ppm, and 0.1 to 3 ppm, respectively. These metasediments have ΣREE contents varying from 15.5 to 30.1 ppm, LREE contents from 11 to 17.9 ppm, HREE from 4.4 to 11.8 ppm, and slightly negative to positive Eu/Eu* anomalies ranging from 0.3 to 1.3, suggesting vigorous changes in the physic-chemical conditions.

5.4.2. Classifications and Origin of Schists

Na₂O-FeO + MgO-K₂O ternary diagram [64] was used to determine the origin of metasedimentary samples. Plotting of the investigated metasediments on this diagram revealed that they fall in the greywacke field, which suggests a sedimentary origin (Figure 12a). The FeO^t-MgO-Al₂O₃ ternary diagram of [65] supports the metamorphic origin of these rocks (Figure 12b). According to the (SiO₂/Al₂O₃)–(K₂O/Na₂O) binary diagram of [66], the metasediments plot in the pelitic greywackes field (Figure 12c). Based on the log (FeO/K₂O) versus log (SiO₂/Al₂O₃) of [67], the samples fall within the Fe-shale field, indicating high content of iron oxides (Figure 12d). K₂O versus Na₂O of [68] divided the greywacke into quartz rich, quartz intermediate, and poor quartz varieties; the studied samples are fitted with quartz intermediate and poor fields. (Figure 12e). Chondrite normalized rare Earth elemental (REE) patterns of the metasediments are shown in Figure 13. The metasedimentary samples are characterized by a moderate degree of REE fractionation, as revealed by their contents (La/Yb)_N ratios (1.2 to 8.3 ppm). The degree of LREE fractionation is quite low, and (La/Sm)_N content ranges from 1.0 to 3.6 ppm, while the heavy REEs are weakly fractionated and (Gd/Yb)_N content range from 0.9 to 5.3 ppm.

5.5. Classifications and Magma Type of Island-Arc Metavolcanic

According to Middlemost’s Na₂O + K₂O vs. SiO₂ diagram [69], metavolcanic samples fall in the basaltic–andesite field (Figure 14a). On the variation diagram of Na₂O + K₂O vs. SiO₂ [60], all samples plot in the subalkaline magma field (Figure 14b). The (AFM) Alkalis-FeO^t-MgO diagram proposed by [60] can discriminate tholeiitic and calc–alkaline affinity. In this diagram, the samples plot within the calc–alkaline field (Figure 14c). On the SiO₂ versus FeO^t/MgO variation diagram of [70], the metavolcanic samples plot in the calc–alkaline field (Figure 14d). The investigated metavolcanic in chemistry resembles Shadli island-arc metavolcanic [71].

5.6. Tectonic Setting and Petrogenesis of the Island-Arc Metavolcanic Rocks

According to Nb versus SiO₂ variation diagram of [72] to discriminate between the active continental margin, orogenic volcanic terrains and immature island arcs. All metavolcanic samples plot in the immature island arcs (Figure 14e). Based on the Th/Yb versus Nb/Yb variation diagram of [73], all samples plot in the oceanic arc field (Figure 14f), suggesting an arc-related (island-arc) setting [72].

5.7. Tectonic Model and Emplacement Mechanism

Egypt’s Central Eastern Desert (CED) is divided into two ophiolite belts with controversial genetic systems. The first belt is called Ghadir–Mubarak–Barramiya and is located in the south [26], while the second belt is the Wadi Semna–Fawakhir–Um Gheig belt in the north of CED [21]. The study area is located in the northern belt of CED (Figure 15 and Figure 16), which is affected by NW obduction-related thrust faults resulting from the structure evolution of the Najd fault system (NFS).

The two Cryogenian ophiolitic belts in the CED and the accretionary history of the Eastern Desert of Egypt were overprinted by the Najd fault system (NFS) with fore-arc and back-arc geochemical characteristics (Figure 15a); Abd El-Rahman et al. propose that an intra-oceanic island connected to a NE-dipping subduction zone formed during the tectonic history of the CED [4]. A fore-arc ophiolite belt, however, stretches from El Barramiya [26] in the southern portion of the CED (SCED) to Fawakhir in the northern section of the CED (NCED). In addition, a back-arc basin parallel to the fore-arc belt formed to the east (in modern coordinates), and the back-arc basin was closed by a WSW-dipping subduction zone [21] (Figure 15b).

Figure 15. The controversial models for the evolution of the CED ophiolite belts according to [21]; (a) Model I after [74]; (b) Model II after [4] showing the distribution of serpentinite in the Eastern Desert of Egypt. Zones with blue colour in both models are the fore-arc association, while the zones with red colour are the back-arc association.

Figure 15. The controversial models for the evolution of the CED ophiolite belts according to [21]; (a) Model I after [74]; (b) Model II after [4] showing the distribution of serpentinite in the Eastern Desert of Egypt. Zones with blue colour in both models are the fore-arc association, while the zones with red colour are the back-arc association.

Alternatively, some studies argue that the Eastern belt (toward the Southern Eastern Desert) is a fore-arc, and the Western belt (toward the Northern Eastern Desert) is a back-arc, both of which have opposing directions. This theory is more consistent with the overall structural patterns (Figure 16a). The metasomatized serpentinites of Kab Amiri match the majority of Egyptian ophiolites’ proposed tectonic settings, which featured oceanic lithosphere fragments deposited over a subduction zone in a fore-arc setting, according to the aforementioned data (Figure 16b) [75,76,77].

In the suggested models for the CED’s development, the presence of the paired belts is indicated by ophiolite belts. Having geochemical affinities in the fore-arc and back-arc [4]. The first piece of evidence for this concept is the age of the fore-arc belt, which should be older than the back-arc belt. However, the beginning of a back-arc basin happens later to the spreading at the front of the subduction zone at the fore-arc setting [21].

Figure 16. (a) The proposed model of Neoproterozoic fore-arc ophiolite belts in the northern Arabian-Nubian Shield at Central Eastern Desert (CED), Egypt [21,74]. (b) Tectonic model for the evolution of the studied ultramafic rocks and the Arabian–Nubian [76,77,78].

Figure 16. (a) The proposed model of Neoproterozoic fore-arc ophiolite belts in the northern Arabian-Nubian Shield at Central Eastern Desert (CED), Egypt [21,74]. (b) Tectonic model for the evolution of the studied ultramafic rocks and the Arabian–Nubian [76,77,78].

6. Conclusions

• Most previous studies mainly concern granitic rocks’ geology, petrography, and chemistry. We carried out detailed field, petrographical and geochemical studies on the serpentinites, spilite, metasediments and metavolcanic (metabasalt) at Wadi Kab Amiri.
• Dismembered ophiolitic rocks are distributed in the northern and eastern parts of the study area and comprise serpentinites, talc carbonates and metabasalt rocks. Serpentinites in the present area are the oldest rock unit. Petrographically, the serpentinites are composed of serpentine minerals (90%), and olivine with minor amounts of carbonates and opaques. Geochemically, the serpentinites resemble those of depleted mantle peridotites (harzburgite-dunite) and tend to have abyssal geochemical characteristics.
• Kab Amiri metabasalt is related to the older metavolcanic (OMV), characterized by large exposure of pillow structures. They consist of plagioclase, chlorite, carbonate and epidote; they are also characterized by cryptocrystalline grain size and amygdaloidal texture. The ophiolitic metabasalt are sub-alkali basalt affinities, tholeiitic and formed in the fore-arc basin. The chondrite normalized rare Earth elemental (REEs) patterns of the metabasalts show low fractionated patterns.
• Island arc assemblages are represented by metasediment and metavolcanic rocks. Metasediments rocks comprise quartz–actinolite–chlorite schist and slate. They are mostly foliated and highly folded. Microscopically, the schists are essentially composed of amphibole and mica minerals, plagioclase, and quartz. In contrast, slate samples are composed mainly of quartz, sericite, and chlorite and are related to oceanic island arc tectonic setting.
• Large exposures of andesitic basalt in the southern and northeast parts of the mapped area represent metavolcanic rocks. Microscopically, the island arc metavolcanic is classified as metabasalt (plagioclase and mafic minerals) and metapyroclastic (porphyroblasts of clastic plagioclase, quartz, carbonate and perthite embedded in a fine-grained groundmass of chlorite and epidote). Geochemically, the metavolcanic are basaltic-andesite andesite in composition, and calc-alkaline affinity is related to the island arc tectonic setting. Chondrite-normalized rare earth element (REE) patterns show slightly fractionated patterns.
• The large serpentinite bodies are concentrated on the CED’s western side, extending in an NNW–SSE direction. To the west of the arc-back-arc assemblages, arc-forearc assemblages are still visible, proving that the intra-oceanic island arc system was formed over an E-dipping subduction zone (present coordinates). Both assemblages define a suture zone in the CED by extending as NW-SE belts.
• The geochemical signature of subduction increases in the direction of the south in arc-back-arc assemblages. Bimodal volcanism’s prevalence in the south suggests that a back-arc basin has been gradually expanding from the north to the south, but the rift island arc assemblages farther south have been preserved.