Introduction

Peridotites as ultramafic mantle section cropped out at the surface of the Earth can offer important information on the geochemical and geodynamic evolution of the upper mantle (e.g., Sano and Kimura 2007; Dai et al. 2011, 2013; Uysal et al. 2012; Khedr et al. 2014).

The Arabian–Nubian Shield (ANS) marks the northern part of the East African Orogen and was formed close to the end of the Proterozoic due to collision between major fragments of East and West Gondwana and termination of the Mozambique Ocean (Stern 1994; Johnson et al. 2003). Accretion of intra-oceanic island arcs and collision between these arcs and a continental margin produced the ANS and created suture zones outlined by ophiolites (Stern 1994, 1996; Stern et al. 2004). Ophiolite sequences are widely distributed in the central and southern segments of the Eastern Desert (ED) of Egypt (Fig. 1). The complete ophiolite successions (e.g., Wadi Ghadir, Gabal Gerf and Fawakhir) are scarce, whereas the dismembered ophiolitic fragments are several (El-Sharkawy and El-Bayoumi 1979; El-Bayoumi 1980; Shackleton et al. 1980; El-Sayed et al. 1999; Abd El-Rahman et al. 2009a, b; El Bahariya 2008, 2012; Abdel-Karim et al. 2016). Fawakhir ophiolitic complex dates 736.5 ± 1.2 Ma (Andresen et al. 2009).

Fig. 1
figure 1

Ophiolites distribution in the central and southern Eastern Desert of Egypt and the location of Um Halham (H) and Fawakhir (F) ophiolites (after Shackleton 1994). Inset displays the general map of Egypt and the location of figure is shown. CED Central Eastern Desert, SED Southern Eastern Desert

Full size image

A SSZ setting of the ED ophiolites is commonly acknowledged (El Bahariya and Arai 2003; Azer and Khalil 2005; Azer and Stern 2007; Ahmed et al. 2012; Abdel-Karim et al. 2016, 2017). However, debate exists on whether they were formed in back-arc or fore-arc settings. A back-arc setting is usually concluded based on the geochemical signatures of the volcanic units of the ophiolite assemblage (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. 1996, 2008). In contrast, based on the mineral and bulk-rock data of the serpentinites a fore-arc setting was proposed (Stern 2004; Azer and Stern 2007; Abd El-Rahman et al. 2012; Azer et al. 2013; Abdel-Karim et al. 2015, 2016, 2017). To contribute toward solving this issue; we present bulk-rock and mineral compositional data for the metaultramafics of Um Halham and Fawakhir areas to better constrain their geochemical and geodynamic evolution.

Field characteristics

Um Halham

Um Halham area (Fig. 2) is situated between latitudes 26°19′48″–26°23′24″ N and longitudes 33°41′24″–33°48′36″ E in the western part of the CED and is covered mainly by volcano-sedimentary sequence thrusted to south over El-Rubshi ophiolites. The volcano-sedimentary sequence around Um Halham represents an interbedded sequence of clastic metasediments and metavolcanics. The western part of the sequence starts at the base with clastic metasediments enclosing thin intercalations of iron ore bands, up to 10 cm thick, interbedded with metavolcanics (Abu El-Ela 1990). The eastern part of the sequence is represented mainly by pyroclasts and minor lava flows interbedded with metasediments. Um Halham ophiolites include talc–carbonates, serpentinites and chromitites. The area is intruded by small bodies of syn-tectonic granites and Abu Hayaya post-tectonic granites (Fig. 2). Serpentinite forms mountainous ridge with sharp irregular peaks (Fig. 3a). Talc–carbonates are thrusted over the volcano-sedimentary sequence (Fig. 3b). In some places, talc–carbonates exhibit conspicuous cavernous structure (Fig. 3c). Sometimes, chromite-rich serpentinites occur as fragments of variable sizes and shapes within talc–carbonates (Fig. 3d).

Fig. 2
figure 2

Geologic map of Um Halham area (after Abu El-Ela 1990). Sample numbers are indicated

Full size image
Fig. 3
figure 3

Field photographs of Um Halham showing a serpentinites form mountainous-ridge with sharp irregular peaks; b talc–carbonates thrusted over volcano-sedimentary sequence; c talc–carbonates exhibit conspicuous cavernous structure; d chromite-rich serpentinite fragments in talc–carbonate rocks

Full size image

Fawakhir

Fawakhir area (Fig. 4) is located 93 km west of the Red Sea coast along Qift–Quseir asphaltic road between latitudes 25°57′36″–26°01′48″ N and longitudes 33°34′12″–33°40′12″ E. The total mapped area covers ~ 88 km2. Fawakhir ophiolites situated between the Dokhan volcanics to the west and the Meatiq Dome ~ 15 km to the east. East dipping thrust zone characterized the western contact between the Fawakhir ophiolite suite and the Hammamat sediments, whereas its eastern contact with the Meatiq rocks is characterized by a zone of tectonic mélange, mylonitized rocks and highly-deformed ‘flaser’ gabbro. They contain mostly serpentinites, metagabbros, and metabasalts (El-Sayed et al. 1999). The Dokhan Volcanic rocks constitute a small mass located at the extreme western part of the study area. The serpentinites and metagabbros are intruded by Fawakhir post-tectonic granites with sharp contact (Fig. 5a). The serpentinites bordered Fawakhir granitoids from the west, north and south, however most of the eastern side lacks them (Fig. 4). Many shear zones cut across the serpentinized ultramafics which are altered to talc–carbonates along these shears (Abd El-Rahman et al. 2009a). Sometimes, the massive serpentinites show talc–carbonate alteration (Fig. 5b). The metaultramafics and adjacent rocks display sharp and distinct contacts. Serpentinites display sharp contact with the underlying mélange. This sharp contact is characterized by a NNW–SSE trending deep thrust fault and a relatively narrow dark green schistose amphibolite band between the ultramafic and the mélange (Hassanen 1985). The adjacent country rocks lack the thermal metamorphic effect of the serpentinites indicating tectonic emplacement of the serpentinites (El-Sayed et al. 1999).

Fig. 4
figure 4

Geologic map of Fawakhir area (after El-Sayed et al. 1999)

Full size image
Fig. 5
figure 5

Field photographs of Fawakhir area showing a sequence of serpentinite, metagabbro and post-tectonic granite along Qift–Quseir Road. Photo looking east; b massive serpentinite altered to talc–carbonate (buff color)

Full size image

Petrography

The studied ophiolitic rocks are partially to completely serpentinized and comprise both massive and sheared varieties. Serpentinized peridotites and serpentinites belong to massive varieties while talc– and quartz–carbonates affiliated to sheared serpentinites. Both the Um Halham and Fawakhir metaultamafic rocks comprise serpentinized peridotites, serpentinites together with talc– and quartz–carbonates. They contain olivine, chromite and pyroxene relicts. Chromites are represented by homogenous and zoned crystals. The zoned chromites have darker cores compared to the outer rims which have lighter gray color and higher reflectance.

Serpentinized peridotites in Um Halham contain relicts of olivine (Fig. 6a) and clinopyroxene (Fig. 6b). Serpentine, tremolite–actinolite, chlorite and talc are the main metamorphic products. As metamorphism progresses, serpentinized peridotites change to serpentinites and carbonate–serpentinites. Serpentinites contain mainly serpentine minerals together with minor amounts of carbonates and opaques. Olivine relict and talc are recorded in few samples. Serpentinites show both pseudomorphic and interpenetrating textures. Serpentines formed after olivines exhibit mesh texture (Fig. 6c), whereas those developed after orthopyroxene display bastite texture (Fig. 6d). Sometimes, bastite texture accompanies schiller structure where magnetite grains outline original orthopyroxene cleavage planes (Fig. 6d). Carbonates exist as fine- to medium-grained patches substituting bastites or as veinlets substituting pre-existing chrysotile. Talc–carbonates consist principally of clusters or sparse patches of carbonates set in talc matrix, whereas quartz–carbonates essentially contain carbonates together with minor amounts of quartz. Chromites occur as zoned (Fig. 6e) and homogenous crystals with subordinate amounts of magnetites and sulfides.

Fig. 6
figure 6

Photomicrographs taken under crossed nicols (e and j are back-scattered electron; BSE) displaying a fractured olivine crystal altered at margin to talc and serpentine minerals in serpentinized peridotite of Um Halham; b pyroxene crystal replaced by serpentine minerals in serpentinized peridotite of Um Halham; c mesh texture with isotropic mesh center in serpentinite of Um Halham; d bastite texture associated with schiller structure in serpentinite of Um Halham; e zoned chromite crystal mantled by Cr–magnetite with sharp contact in serpentinite of Um Halham; f mesh texture with olivine mesh center in serpentinized peridotite of Fawakhir; g talc and relics of pyroxene in serpentinized peridotite of Fawakhir; h chrysotile veinlet traversing antigorite matrix forming interpenetrating texture in serpentinite of Fawakhir i hourglass texture in serpentinite of Fawakhir; j fractured chromite crystal filled with serpentine minerals in serpentinite of Fawakhir

Full size image

Serpentinized peridotites in Fawakhir area consist of serpentine together with minor relics of olivine, chromite and pyroxene. Tremolite, talc, chlorite and iron oxides are secondary minerals. Serpentine minerals after olivine show pseudomorphic textures. The serpentinization affected olivine grain boundaries and fractures forming olivine mesh center (Fig. 6f). In the serpentinized peridotites pyroxene relicts present and sometimes replaced by talc (Fig. 6g). Serpentinites consist mainly of serpentine minerals associated with variable amounts of talc, carbonate and opaques. They exhibit pseudomorphic, interpenetrating (Fig. 6h) and hourglass textures (Fig. 6i). Talc–carbonates are fine-grained rocks with brownish yellow to reddish brown color. They contain essentially talc and carbonates with rare serpentine, chlorite and opaques. Talc occurs as vein-like and in the groundmass with clusters and patches of carbonate minerals. Zoned and homogenous chromites (Fig. 6j) occur with minor sulfides and magnetites.

Analytical techniques

A total of 46 mineral chemical analyses of olivine (12), clinopyroxene (12) and chromite (22) were performed using Phillips XL30 analyzer at the Central Laboratories of the Geological Survey in Cairo, Egypt. Operating conditions were 15 kv accelerating voltage, 20 nA beam current and 1 µm beam diameter. The raw data were corrected using ZAF correction factors. The analytical results (normalized to 100%) of different minerals are given in Supplementary Table S.

Eighteen representative samples of Um Halham and Fawakhir were analyzed for major and trace elements. Major elements were analyzed using XRF techniques at the Institute of Geothermal Science, Noguchibaru, Beppu city, Japan. Fused beads were prepared from a mixture of specimen and Li2B4O7 flux, diluted 1:5 at 950 °C to achieve accurate and precise results. Replicate analyses provided an overall procedural uncertainty better than 2% for major elements. The trace elements were analyzed after lithium metaborate/tetraborate fusion as well as nitric acid digestion of a 0.2-g sample using inductively coupled plasma–mass spectrometry techniques (ICP–MS) at the Acme Analytical Laboratories in Vancouver, BC Canada. The accuracy and reproducibility of this method lie within 2%. Bulk-rock geochemical analyses with detection limits are presented in Table 1.

Table 1 Bulk-rock major (wt%), trace and rare-earth element (ppm) data of the studied metaultramafics
Full size table

Results

Mineral chemistry

Chromite chemical analyses are presented in Table S. Homogenous chromites and internal cores of zoned crystals are classified as Al–chromites while the outer rims have Cr–magnetite compositions (Fig. 7a). The fresh Al–chromite cores plot close to the Cr–Al join whereas the altered Cr–magnetite rims plot along Cr–Fe3+ join (Fig. 7a). Spinel in mantle peridotites is relatively unaffected by alteration and its Cr# particularly appears to remain unchanged during sub-solidus stage (Arai 1994). Geochemical consequences due to spinel alteration are discussed below. Al–chromites in the studied rocks have higher Cr# (> 60) akin to those of ophiolitic ultramafics of ANS including SED and CED (Fig. 7b, c) (Stern et al. 2004; Farahat et al. 2011; Ahmed et al. 2012; Abdel-Karim et al. 2016; Obeid et al. 2016). Their TiO2 and Al2O3 contents are comparable to SSZ peridotites and overlap those of peridotites from Izu–Mariana and New Caledonia arcs (Fig. 7d; Kamenetsky et al. 2001).

Fig. 7
figure 7

Plot of chromites on a Cr–Fe3+–Al ternary diagram. Greenschist and lower-amphibolite metamorphic facies compositional fields are after Evans and Frost (1975) and Suita and Strieder (1996); b Cr# versus TiO2. The abyssal peridotites compositional field and trends of the effect of MORB–melt reaction on refractory abyssal peridotite spinels and of boninite–melt reaction on refractory SSZ peridotite spinels are from Choi et al. (2008); c Cr# versus Mg# diagram exhibit fields of different tectonic setting (Dick and Bullen 1984; Bloomer et al. 1995; Ohara et al. 2002; Stern et al. 2004). Chromite compositional fields of SED (Ahmed et al. 2012) and CED (Farahat et al. 2011) serpentinites are also shown; d Al2O3–Ti2O. Spinels compositional fields of SSZ and MORB-type peridotites as well as those of fore-arc peridotites from the Izu–Mariana and New Caledonia are after Kamenetsky et al. (2001)

Full size image

Olivines range in composition from Fo85.54–95.30 in Um Halham and from Fo89.33–91.65 in Fawakhir and are like primary olivine relicts of ophiolitic metaultramafics from the ED of Egypt (Fo89–96) (Table S) (Khudeir 1995; Khalil and Azer 2007; Khedr and Arai 2013; Obeid et al. 2016). These compositions correspond to chrysolite and forsterite. Moreover, their Fo-enriched nature is comparable to those observed in the ANS ophiolites and fore-arc peridotites further confirming their primary nature and suggest that the studied rocks are residual mantle after extensive melting (Stern et al. 2004). These inferences are also supported by spinels compositional data of the present study.

Clinopyroxenes have Mg# varying between 0.82 and 0.90 in Um Halham, and between 0.54 and 0.60 in Fawakhir. Their compositions rang from augite (En47.94–65.38Wo27.32–41.27Fs6.46–10.53Ac0.27–0.47) in Um Halham to pigeonite (En44.92–48.35Wo10.48–13.67Fs39.42–43.01Ac0.22–0.53) in Fawakhir (Table S) (Fig. 8a). Most of the clinopyroxenes in the serpentinized peridotites have Al2O3 (> 0.2 wt%) and Cr2O3 (> 0.4 wt%) contents higher than those observed in metamorphic clinopyroxenes reflecting magmatic origin (Nozaka 2010). The analyzed clinopyroxenes have low TiO2 content characteristic of non-alkaline rocks (Le Bas 1962) and display a strong affinity to intra-oceanic fore-arc boninites (Fig. 8b). Their lower Na2O and K2O contents (≤ 0.15 and ≤ 0.04 wt%, respectively) indicate depleted nature of their host rocks. Moreover, the low Na2O and TiO2 suggest sub-oceanic source (Delavari et al. 2009).

Fig. 8
figure 8

Plot of the analyzed pyroxenes on a En–Wo–Fs ternary diagram (after Morimoto et al. 1988); b SiO2–TiO2–Na2O ternary diagram (after Beccaluva et al. 1989)

Full size image

Geochemistry

Bulk-rock chemistry of the studied rocks from Um Halham and Fawakhir is listed in Table 1. To diminish the effect of variable element dilution result from serpentinization process, major element oxides were recalculated on an anhydrous basis and plotted volatile-free. The rocks in both areas display different stages of metamorphism starting from the least affected serpentinized peridotites (LOI = 6.76–11.02) and then serpentinites (LOI = 12.46–26.78) to the most affected talc– and quartz–carbonates (LOI = 14.12–41.11). MgO is less affected by serpentinization and its higher content (MgO = 29.82–48.71 wt%) indicates more depleted mantle nature (Frey et al. 1985; Parkinson and Pearce 1998). The Mg# ranges from 0.85 to 0.94 in Um Halham and from 0.89 to 0.92 in Fawakhir, and are comparable to those from modern oceanic peridotites (Bonatti and Michael 1989) indicating also a limited adjustment of Mg and Fe. The depletion of Na2O (< 0.05 wt%) and K2O (∼ 0.01 wt%) further confirms this implication (e.g., Abdel-Karim et al. 2016). Geochemical consequences due to alteration and metamorphism are also discussed below.

Al2O3 content seems to be relatively not influenced by serpentinization process demonstrating that the bulk-rock Al content typically reflects its original primary concentration (Bonatti and Michael 1989). The studied rocks have Al2O3 contents (0.15–3.52 wt%) overlap those from oceanic and active margin peridotites and fore-arc and Pan-African serpentinites (Floyd 1991; El Bahariya and Arai 2003; Abdel-Karim and Ahmed 2010). Their SiO2/MgO ratios and Al2O3 contents are akin to ophiolitic peridotite (Bodinier and Godard 2003; Abdel-Karim et al. 2016) (Fig. 9a). Serpentinized peridotite of Um Halham (H.21) has higher Al2O3 (3.52 wt%) and CaO (8.18 wt%) compared to that of Fawakhir (F.98) (Al2O3 = 1.54 wt% and CaO = 0.9 wt%). Furthermore, serpentinites in both areas have low Al2O3 contents (0.15–0.86 wt%) and have a wide range of CaO (0.06–8.84 wt%). The low Al2O3 and CaO contents in most samples are comparable to fore-arc peridotites of ED (Zimmer et al. 1995; Azer and Khalil 2005; Azer and Stern 2007; Azer et al. 2013; Abdel-Karim et al. 2016). The Al2O3/SiO2 and MgO/SiO2 ratios are comparable to those of Arabian shield and fore-arc peridotites (Fig. 9b; Parkinson and Pearce 1998; Pearce et al. 2000; Ahmed and Habtoor 2015).

Fig. 9
figure 9

a Bulk-rock Al2O3 versus SiO2/MgO ratio diagram. Fields of ophiolitic peridotite, ophiolitic gabbro and MORB are from Bodinier and Godard (2003). ED field is ophiolitic peridotite compositions from both the Central and Southern Eastern Desert of Egypt (after Azer and Khalil 2005; Azer and Stern 2007; Azer et al. 2013; Abdel-Karim et al. 2016); b MgO/SiO2 versus Al2O3/SiO2 diagram. Primitive and depleted mantle values are from McDonough and Sun (1995) and Salters and Stracke (2004), respectively. The “terrestrial array” represents the bulk silicate Earth evolution (Jagoutz et al. 1979; Hart and Zindler 1986). Fields of abyssal and fore-arc peridotite are from Niu (2004), Pearce et al. (2000) and Parkinson and Pearce (1998). AS represents the field of Arabian Shield ophiolitic peridotite (Ahmed and Habtoor 2015)

Full size image

The studied rocks are enriched in compatible trace elements (Cr = 1338–3082 ppm, Ni = 557.6–2882.2 ppm and Co = 73.3–117.8 ppm) reflecting development from a depleted mantle peridotite source. Their trace element patterns (Fig. 10a) show fluid mobile element (FME; such as Cs, U, and Pb) enrichments resembling those of fore-arc serpentinites along with Sr enrichment which is not observed in serpentinized peridotite (Pearce et al. 2000; Song et al. 2009). These enrichments in FME, most likely point to the higher concentrations of these elements in the hydrating fluids of the mantle wedge. These hydrating fluids are rich in Cs, Pb, and Sr indicating that they were produced from sediment dehydration during shallow subduction (e.g., below fore-arcs; Deschamps et al. 2013 and references therein). Moreover, the high LILE contents [e.g., Ba/Th (up to 350), Ba/La (up to 70)] overlap those observed on mantle wedge serpentinites from diverse fore-arc systems suggesting interaction with sediment-derived fluids (Aziz et al. 2011; Kodolányi et al. 2012; Deschamps et al. 2013). Most of the trace elements (except Cs, U, Sr and Pb) are depleted compared to primitive mantle values (Fig. 10a; McDonough and Sun 1995) similar to mantle wedge serpentinites (Deschamps et al. 2013). Both the positive Pb-anomaly and the negative Nb-anomaly characterize the studied rocks (Fig. 10a). The overall low concentrations of trace element suggest high degrees of melt extraction of the mantle protolith, similar to mantle wedge serpentinites (Deschamps et al. 2013).

Fig. 10
figure 10

Data sources: fore-arc peridotites (Pearce et al. 2000; Song et al. 2009); SED ophiolitic ultramafics (Abdel-Karim et al. 2016)

a Primitive mantle-normalized trace element patterns; b chondrite-normalized REE patterns. Normalization values are after McDonough and Sun (1995).

Full size image

The REEs mobility during serpentinization process is insignificant, so they preserve the geochemical characteristics of the original mantle protolith (Deschamps et al. 2013). Bulk-rock REE contents of the studied rocks are relatively variable. Serpentinized peridotites have LREE composition ranging from ~ 2 to ~ 5 Chondrite and HREE from ~ 0.4 to ~ 2 Chondrite, whereas serpentinites have LREE composition ranging from ~ 1.5 to ~ 5 Chondrite and HREE from ~ 0.08 to ~ 0.3 Chondrite (Fig. 10b). Moreover, talc– and quartz–carbonates have LREE compositions varying from ~ 1.5 to ~ 4 Chondrite and HREE from ~ 0.2 to ~ 2 Chondrite (Fig. 10b). The studied rocks show no correlations between REEs and LOI confirming that the REE signatures of the mantle protolith were preserved during serpentinization and metamorphism (e.g., Savov et al. 2005a, b; Deschamps et al. 2013).

Serpentinized peridotite of Um Halham has both enriched LREE [(La/Yb)N = 2.1] and HREE [(Gd/Yb)N = 1.03] patterns as do quartz–carbonate [(La/Yb)N = 2.4 and (Gd/Yb)N = 1.32] (Fig. 10b). On the other hand, serpentinites and talc–carbonate rocks exhibit enriched LREE patterns with (La/Yb)N = 3.1–32.3 and depleted HREE patterns with (Gd/Yb)N = 0.47–0.83. Generally, serpentinites represent the most LREE-enriched rocks. Serpentinites, talc– and quartz–carbonates show variable positive Eu-anomaly compared with serpentinized peridotite (Fig. 10b).

Serpentinized peridotite of Fawakhir is LREE-enriched [(La/Yb)N = 16.74] relative to HREE [(Gd/Yb)N = 2.21] patterns (Fig. 10b). Moreover, serpentinites and quartz–carbonate have LREE patterns [(La/Yb)N = 3.6–14.35] higher than HREE patterns [(Gd/Yb)N = 0.55–2.32]. On the other hand, talc–carbonate (F.117) exhibits both LREE- and HREE-enriched patterns [(La/Yb)N = 11.5 and (Gd/Yb)N = 17.5]. In contrast to serpentinized peridotites, the analyzed samples from Fawakhir also display variable positive Eu-anomaly (Fig. 10b).

Discussion and conclusion

Geochemical consequences due to alteration and serpentinization

Generally, the ophiolitic ultramafics in the ED of Egypt are affected by low-grade green schist to medium-grade amphibolite facies metamorphism and altered to serpentinite and/or to combinations of serpentine, talc, chlorite, carbonates and magnetite (e.g., El-Sayed et al. 1999; Farahat 2008; Khedr and Arai 2013). Carbonate alteration affected the Egyptian ultramafic rocks but its timing and source remain controversial. It was attributed to mixing between CO2-rich (mantle-derived) fluids and remobilized sedimentary carbonate (Stern and Gwinn 1990). Pure mantle (CO2-bearing) source is also proposed based on stable isotope (i.e. O, C) (Boskabadi et al. 2017). Moreover, some magnesite veins in serpentinites from the ED were ascribed to an influx of mantle and metamorphic-degassing of CO2 (e.g., Ghoneim et al. 2003; Hamdy and Lebda 2007).

Despite the mineralogical modifications in peridotites during serpentinization processes, geochemical studies of serpentinites show insignificant changes in major elements (except for Ca) at the hand-specimen scale (e.g., Mével 2003; Deschamps et al. 2010, 2013). However, except for two samples (H.21 and H.28), the low CaO contents (0.06–4.63 wt%) in the serpentinized peridotite and serpentinites reflect limited effect of Ca–metasomatism which is further supported by the absence of appreciated correlation between CaO and LOI in all samples. Moreover, bulk-rock data of serpentinized peridotites similarly indicate insignificant modifications occur in the trace element compositions (excluding U or Sr) during serpentinization (e.g., Niu 2004; Paulick et al. 2006; Deschamps et al. 2013). Consequently, major and trace element compositions can be used to define the nature of the serpentinites protolith in subduction zones (Chalot-Prat et al. 2003; Hattori and Guillot 2007; Deschamps et al. 2013).

Low-temperature serpentinization process leads to initial growth of slightly Fe-enriched chromite because of fluid infiltration along chromite cracks and grain boundaries (Barnes 2000). At higher temperature, further fluid access and reaction give rise to extensive magnetite replacement of chromite (Barnes 2000). Chromite alteration caused by serpentinization process in the late stage of magmatism and possibly accompanying emplacement of the ophiolite complex (e.g., Khudeir et al. 1992; Khalil and Azer 2007). The altered Cr–magnetite rims have higher total iron and lower Cr2O3, Al2O3, MgO and Cr# than the fresh Al–chromite cores (Table S) indicating an alteration event. Moreover, the Cr–magnetite plot on the Cr–Fe3+ joins close to Fe3+ apex (Fig. 7a) reflecting Fe2O3 increase and loss in Cr2O3 and Al2O3 during alteration. Their high Fe3+ reflects oxidation conditions during alteration (Anzil et al. 2012).

Metamorphic grade

Chromite cores continually equilibrated with magnetite rims document metamorphic grade conditions (Barnes 2000). The relative proportions of the trivalent ions (i.e. Cr3+, Al3+ and Fe3+) of chromite are unaffected by metamorphism up to lower temperature amphibolite facies implying restricted mobility of these elements occurred under lower amphibolite (Barnes 2000). Therefore, chromite in lower temperature amphibolite facies preserves its primary igneous chemistry and can be used to estimate the metamorphic grade (Barnes 2000). According to Fig. 11, almost all the Al–chromite cores are equilibrated at temperature below ~ 500–550 °C corresponding to lowest amphibolite facies metamorphism (Sack and Ghiorso 1991; Barnes 2000). So, they reflect magmatic composition not influenced by metamorphism (e.g., Barnes 2000). On the other hand, altered chromite rims have nearly pure magnetite compositions with restricted Cr-solubility indicating magnetite development at < 500 °C (Fig. 11; Sack and Ghiorso 1991; Barnes 2000).

Fig. 11
figure 11

Spinel data from the studied rocks compared with Sack and Ghiorso spinel stability fields for chromite and magnetite (after Barnes 2000)

Full size image

Indeed, this inference is confirmed by plotting the analyzed Al–chromite cores in the primary spinel fields of greenschist and lower-amphibolite metamorphic facies (Fig. 7a: Evans and Frost 1975; Suita and Strieder 1996). Thus, the studied rocks experienced metamorphism grade from greenschist up to lower-amphibolite metamorphic facies (Figs. 7a, 11).

Modification of the mantle source by subducted components

The studied rocks exhibit very low Nb, Ta, Zr and Hf concentrations along with sub-chondritic ratios of Nb/Ta (0.3–16) and Zr/Hf (mostly 1–20) indicating depletion of their mantle source by earlier melting extraction event (e.g., Yang and Zhou 2009).

Mantle source modification by continental materials and/or subduction-related slab melts/fluids has been confirmed for ophiolitic utramafics (e.g., Sharma and Wasserburg 1996; Gruau et al. 1998; Delavari et al. 2009; Ahmed et al. 2012; Abdel-Karim et al. 2016). Ahmed et al. (2012) which suggested that Gebel Mudarjaj ultramafics were formed by contamination of arc-related magmas by lower crustal gabbroic rocks. Recently, Abdel-Karim et al. (2016) indicated that the Gerf serpentinites could have been generated by contamination of their mantle source with crustal material and/or subduction-related slab fluids.

On the (La/Sm)N versus (1/Sm)N diagram of Sharma and Wasserburg (1996) (Fig. 12), a comparison presented between the studied metaultramafics and model peridotites produced from contaminated residual harzburgites with MORB melts or continental crust components. The studied rocks display LREE-enrichment [(La/Sm)N average 10; Fig. 12]. Moreover, they plot around lines demonstrating mantle modified with crustal materials and/or subduction-related slab melts/fluids (Fig. 12; Gruau et al. 1998). However, their low Th/Nb ratios (mostly 0.1–2) and negative Zr anomalies argue against significant crustal contamination (e.g., Yang and Zhou 2009). Thus, their high (La/Sm)N resulted from subduction-related slab melts/fluids influx during the evolution of the mantle in a subduction zone setting. The low Ce/Pb (average 1.6) and high Ba/Nb (average 191) together with the negative Nb anomalies further support the input of subduction-related slab melts/fluids into their mantle source. The origin of the LREE-enriched Trinity ophiolite (Fig. 12) has also been attributed to contamination of their mantle source (mantle restite) with crustal components (Gruau et al. 1998). This contamination could occur during evolution of the mantle in SSZ setting or following ophiolite complex obduction on the continental crust (Gruau et al. 1998; Delavari et al. 2009). The trend of increasing (La/Sm)N indicates variable input of subduction-related slab melts/fluids into the mantle source of the ED ophiolites (Fig. 12).

Fig. 12
figure 12

Sharma and Wasserburg (1996) diagram showing chondrite-normalized (La/Sm)N versus (1/Sm)N. MORB mid-ocean ridge basalts, UM upper mantle composition, UDM ultra-depleted melt composition, CC continental crust composition, HZ1, HZ2 and HZ3 model harzburgite compositions. Fields of SED and Trinity ophiolite are shown for comparison (Gruau et al. 1998; Abdel-Karim et al. 2016)

Full size image

Protolith geochemical fingerprints

Several studies have shown insignificant mobility of major elements and protolith geochemical fingerprints were preserved during serpentinization process (e.g., O’Hanley 1996; Mével 2003; Niu 2004; Deschamps et al. 2013).

The low CaO contents in serpentinized peridotite and serpentinites are comparable to values noticed on ophiolitic peridotites (Bodinier and Godard 2003), whereas the two samples (H.21 and H.28) with high CaO contents are like serpentinized lherzolites (Deschamps et al. 2013). Moreover, the low Al2O3/SiO2 ratios (mostly < 0.03) of serpentinized peridotite and serpentinites akin to fore-arc mantle wedge serpentinites and indicate that their protoliths had underwent partial melting before serpentinization process which has no influence on this ratio (e.g., Snow and Dick 1995; Paulick et al. 2006; Deschamps et al. 2013). Also, their low MgO/SiO2 ratios (< 1.1) resemble serpentinized lherzolite and harzburgite (Deschamps et al. 2013). They have low TiO2 contents (0.0–0.08 wt%) compared to depleted mantle composition but like subduction zone serpentinites (Salters and Stracke 2004; Deschamps et al. 2013).

Trace element compositions (principally REEs) of serpentinized mantle rocks can be used to interpret the nature of the mantle protolith and the interactions with fluid/melt (Deschamps et al. 2013). The low HFSE content reveals high degrees of melt extraction (e.g., Parkinson and Pearce 1998). Positive Pb-anomaly present on spider-diagrams may reflect a protolith nature or fluid percolation during serpentinization (Fig. 10a) (Godard et al. 2008; Deschamps et al. 2013).

Magma nature

The low Al2O3 content in the studied rocks is consistent with a depleted upper mantle source (Bonatti and Michael 1989). They display higher Mg# numbers together with Cr and Ni enrichment reflecting development from a depleted mantle peridotite source (e.g., Khalil 2007; Abdel-Karim et al. 2016). Their high bulk-rock MgO and SiO2 and low Al2O3 and TiO2, together with clinopyroxenes data suggest boninitic and fore-arc affinity. They have Al–chromites with higher Cr# indicating that they represent highly depleted mantle residues after higher degrees of melting. The degrees of melting were estimated based on fresh chromite chemistry using the empirical equation of Hellebrand et al. (2001) [i.e. Melting degree (F %) = 10*ln (Cr#) + 24]. According to this equation, the degrees of melting vary between 19 and 24% similar to that of Gebel Mudarjaj in the SED of Egypt (Ahmed et al. 2012) and are totally within the range (15–40%) of SSZ peridotites (Pearce et al. 2000; Mellini et al. 2005).

On the TiO2–Cr# diagram (Fig. 7b), the analyzed spinels plot around the boninite–melt reaction trend which is completely different from the MORB–melt trend of the abyssal peridotites. Moreover, the Al–chromites have Cr# and TiO2 compositional features typical of primary spinels from fore-arc depleted mantle peridotites with boninitic characteristics (Dick and Bullen 1984; Arai 1992) (Fig. 7b). The high Cr# (> 60) and Mg# (49–82) of Al–chromite cores are also comparable to those from mantle-derived peridotites which support the primary nature and boninitic affiliation (Fig. 7c) (Roeder 1994; Mondal et al. 2001). Furthermore, the clinopyroxene compositions correspond to that from boninites (Fig. 8b).

Tectonic implications

Serpentinized peridotite and serpentinites have low MgO/SiO2 ratios (< 1.1) according with SSZ peridotites from fore-arc setting (Fig. 9b; Parkinson and Pearce 1998; Pearce et al. 2000). Generally, the studied rocks are depleted in Al2O3 content indicating a SSZ setting (Bonatti and Michael 1989). The negative Nb-anomaly is consistent with SSZ geochemical fingerprints and is typical of fore-arc peridotites (Fig. 10a) (Hawkins 2003; Song et al. 2009). Furthermore, the characteristic positive Pb-anomaly displayed on spider-diagrams is similar to fore-arc peridotites (Fig. 10a) (Hawkins 2003; Deschamps et al. 2013).

Fresh chromite core composition is a good proxy to define the tectonic setting of ultramafic mantle rocks (Dick and Bullen 1984; Arai 1992). The low contents of Al2O3 and TiO2 in the analyzed Al–chromites are identical to peridotites from SSZ setting (Fig. 7d). There is some overlap with the spinels compositional range from the Izu–Mariana fore-arc and New Caledonia mantle peridotites (Fig. 7d). Subduction-related mantle rocks including mantle wedge and fore-arc rocks have spinels with elevated Cr# (> 0.40) (Dick and Bullen 1984; Ozawa 1994; Parkinson and Arculus 1999) which is the case of the studied rocks. The high Cr# (> 60) of the analyzed Al–chromites is comparable to those of modern fore-arc peridotites but clearly higher than spinels from MOR and back-arc peridotites (Fig. 7c; Ahmed et al. 2001; Stern et al. 2004; Farahat et al. 2011; Abdel-Karim et al. 2016). Moreover, the high Cr# is also similar to that observed in the chromitites of boninitic affinity existed in the deeper portion of the mantle section (Rollinson and Adetunji 2015). The Cr# and Mg# relationship shows that the studied rocks are similar to depleted fore-arc peridotites (Fig. 7c) (Ishii et al. 1992; Parkinson and Pearce 1998). Furthermore, they are similar to fore-arc serpentinites from Mariana in the western Pacific and Tso Morari in the NW Himalaya (Ishii et al. 1992; Guillot et al. 2001).