Unravelling the genesis and depositional setting of Neoproterozoic banded iron formation from central Eastern Desert, Egypt

Abstract

The Neoproterozoic banded iron formations (BIFs) are widely occurred in the Egyptian Eastern Desert. This study integrates field observations, petrographic studies, geochemical data, and lead isotopes to construct the genesis and depositional environment of Wadi El-Mis hama BIF deposits. The iron layers, primarily of oxide facies within a volcano-sedimentary sequence, comprise magnetite-rich beds alternating with jaspilite or silicate laminae. The studied BIFs exhibit a dominant composition of SiO₂ and Fe₂O₃^t with relatively low contents of TiO₂ and Al₂O₃. The positive correlation of REEs (La, Sm, Yb) with Zr and low concentrations of HFSEs (Ta, Nb, Th, Hf) indicate a primary formation mechanism of chemical precipitation, maintaining original geochemical signatures. Geochemical patterns show depletion in LREEs, enrichment in HREEs (La/Yb_PAAS = 0.08–0.12), and positive La anomalies (La/La_PAAS = 1.15–8.57), consistent with seawater influence. Additionally, various geochemical discrimination diagrams supported by elevated super-chondritic Y/Ho values (29.6–38.7), weak positive Eu anomalies, and low contents of transition metals (Cu and Zn), point to the interaction of low-temperature (<200°C) hydrothermal fluids (bearing Fe and Si) with seawater during the deposition of the BIFs. The lack of significant negative Ce anomalies along with low Ni/Co, U/Th, and Cu/Zn ratios, imply that the iron mineralization was precipitated from dysoxic to oxic conditions. The geochemical and Pb isotopic data suggest that the iron deposits formed in an extensional geodynamic setting (intra-oceanic arc basin environment) due to the subduction of the Mozambique Plate, with signatures closely matching other Precambrian Algoma-type BIFs.

1 Introduction

Banded Iron-Formations (BIFs) are chemical sedimentary rocks characterized by finely laminated to thinly bedded layers, with alternating dark iron-rich and light silica-rich bands of various thicknesses, indicative of the changing oxidation state of the ocean and atmosphere (Frei et al., 2013). Although banding is a common feature in BIFs, it is not always completely preserved due to varying degrees of metamorphism and deformation (Bekker et al., 2010). BIFs, forming the world’s predominant iron resource (over 95%) due to their global distribution, were mainly precipitated throughout the Archean and Paleo-Meso Proterozoic (∼3.8 and 1.1 Ga) (Pirajno and Yu, 2021; Zhou et al., 2024), with subordinate BIFs deposited from the late Neoproterozoic (Cachoeirinha group, Borborema Province, NE Brazil) (Usma et al., 2021) to early Paleozoic (Taxkorgan Terrane, West Kunlun, China) (Ding et al., 2021). In comparison to the Archean and Paleoproterozoic BIFs, Neoproterozoic BIFs have simple geochemistry and are dominated by oxide facies, and were created in a more oxygenated environment (Cox et al., 2013) (Supplementary Table S1).

Based on the depositional environment, tectonic setting, and size, BIFs are commonly subdivided into three main types, namely Algoma-type, Superior-type, and Rapitan-type (Gross, 1980). The Algoma-type iron formations are comparatively small, stratigraphically associated with submarine volcanogenic successions (felsic to ultramafic rocks), and were likely deposited within extensional geodynamic regimes (deep-water conditions in association with hydrothermal fluids), particularly near spreading centers or in restricted basins (rifted island arcs and back-arc basins). They are characterized by high magnetite content and rapid changes in thickness (Supplementary Table S1). In some cases, they are associated with volcanogenic massive sulfide (VMS) deposits (Bekker et al., 2010). In contrast, the Superior-type is hundreds of meters thick and commonly associated with metasedimentary sequences (i.e., dolostone and shale) without direct relationships with volcanic rocks (Simonson, 1985). Superior-type BIFs were formed in relatively shallow marine conditions over large areas on continental shelves and slopes of passive margins or in intracratonic basins (for example, Australia Hamersley and South Africa Transvaal BIFs deposits) (Li et al., 2014; Pirajno and Yu, 2021). A third type is Rapitan-type, which has been linked to global glaciations (Snowball Earth events) and is distinguished by extreme climate changes and significantly higher atmospheric PO₂ (Planavsky et al., 2014). Significant deposits of this type are found in the Rapitan Group in northern Canada and the Urucum district of Brazil (Klein, 2005). Due to these variations from place to place, it has proved difficult to pinpoint a single depositional model that accurately depicts the formation of all BIFs.

Precambrian BIFs reveal important details about the evolutionary processes and the genesis of chemical sedimentary rocks deposited in marine environments (Isley and Abbott, 1999). It is mostly accepted that iron was generally scavenged from the early oceanic crust. The upwelling of ferrous iron (Fe²⁺) from submarine hydrothermal vents, followed by oxidation into ferric iron (Fe³⁺) in the upper levels and re-deposited on the ocean floor, is linked to most BIFs types (Polat and Frei, 2005). In addition, some recent studies (Li et al., 2015) indicated that continental Fe recycled by microbial dissimilatory iron reduction processes has also been proposed as an important source of Fe in BIFs. Globally, most BIFs have undergone significant diagenetic and metamorphic modifications after their formation (Klein and Beukes, 1993).

The Arabian–Nubian Shield (ANS) hosts several occurrences of BIFs (∼fifteen localities, for example, Wadi Sawawin area in Saudi Arabia, Wadi El-Dabbah area in Egypt, Tambien Group in Ethiopia) (El-Habaak, 2021), which formed during the late Proterozoic and are classified as Algoma-type BIFs. Several small to moderate-size BIFs are recorded and preserved in Egyptian basement rocks (Figure 1; the northern part of the ANS) and accessible today for investigation and study, including: (a) Neoproterozoic BIFs in the Central Eastern Desert (Abu Marwat, Abu Rakib, Abu Diwan El Imra, Gebel El Hadid, El Hundusi, Gebel Semna, Fatira, Um Anab, Um Ghamis El Zarqa, Um Ghamis El Hamra, Um Lassaf, Um Nar, Um Shaddad, Sitra, Wadi El-Dabbah, Wadi Hamama, Wadi Kareim, and Wadi El-Mishama) (Basta et al., 2011; El-Shazly et al., 2019; El-Habaak, 2021) and the South Eastern Desert (scattered areas belonging to the Shadli Metavolcanics Belt) (Khudeir et al., 1988), and (b) Archean-Early Proterozoic BIFs exposed in South Western Desert (Uwaynat area) (Said et al., 1998), and southern Sinai (Wadi Madsus and El Samra area) (Khalid and Oweiss, 1997). Some localities are excluded from the BIF ore type due to the lack of aerial extent and economic potential.

FIGURE 1

The prospective area, Wadi El-Mishama (Figures 2, 3A), lies in the Central part of the Eastern Desert and includes Neoproterozoic BIFs, which were most likely formed during the island-arc stage of the Pan-African Orogeny (Stern and Hedge, 1985; Basta et al., 2011; El-Shazly et al., 2019; El-Habaak, 2021). No geological information or analytical data on the investigated area have yet been published elsewhere. In this contribution, the BIFs and their country rocks from Wadi El-Mishama were investigated. We report the first field observations, petrographic studies, and whole-rock geochemical and Pb isotope data. The aims of the present study are to: (a) constrain the origin of the Wadi El-Mishama BIFs, (b) identify the past geological processes/conditions under which the studied BIFs were precipitated, (c) characterize the Wadi El-Mishama BIFs by comparing them to other representatives Precambrian Algoma-type BIFs in the ANS and around the world, and (d) provide a deeper understanding of the evolution of the Eastern Desert in the Neoproterozoic era. Finally, the data presented here will guide further exploration studies to delineate the iron orebodies.

FIGURE 2

FIGURE 3

2 Geologic setting

2.1 Regional geology

The ANS in the north and the Mozambique Belt in the south are two separate parts of the East African Orogen. It was initiated around 900–870 Ma (Rodinia breakup) and ended around 620 Ma, when east and west Gondwana fragments collided and blocked the Mozambique ocean along the East African–Antarctic Orogen (Ali et al., 2023; Sami et al., 2023a; Sami et al., 2023b). The ANS continued with increasingly intense deformation during the Ediacaran (∼630–550 Ma) as the result of the emplacement of large granitoids, and transtension, transpression, uplift, and exhumation associated with escape tectonics and orogenic collapse (Abd El Monsef et al., 2023; El-Dokouny et al., 2023). The ANS is a collage of high-grade metamorphic core complexes, late Proterozoic juvenile arcs, ophiolitic mélanges, voluminous granitoids, and pre-Neoproterozoic crust enclaves (Sami et al., 2022).

The basement complex of the Egyptian Eastern Desert is dominated by Neoproterozoic Island arcs and ophiolites (Figure 1), which are mostly found in the central and southern regions (Fawzy et al., 2020). The Central Eastern Desert (CED, ∼800 km long) is made up of several thrust sheets, each produced at the base of dismembered ophiolites and capped by island-arc rock assemblages. These juvenile arcs (with variably well-preserved ophiolites) were formed in a subaqueous environment and are characterized by both forearc and back-arc geochemical signatures (Faisal et al., 2020). They are host to a diverse range of economically viable small to medium-sized metallic mineral resources (e.g., BIFs, VMS, orogenic gold). The CED basement rocks suffered regional deformation and metamorphism up to greenschist facies during the period of East African Orogen formation (Abdelfadil et al., 2022). These conditions ranged from 400°C ± 50°C, 4 ± 2 kbar in the northern part of the CED to 520°C ± 30°C, 5 ± 2 kbar in the southern part of the CED (El-Shazly and Khalil, 2016).

The BIFs, which are principally confined between latitudes 25°12′ and 26° 30′N, produced voluminous high-grade iron ores in some localities with ore reserves estimated at ∼17.7, 13.7, and 11.1 Mt for Wadi Kareim, Um Nar, and Um Shaddad BIFs, respectively (El-Habaak, 2021). In terms of geotectonics, three depositional models have been applied to the BIFs in the CED of Egypt: (a) they were formed by chemical precipitation on a continental shelf with the source of iron being continental (El Aref et al., 1993); (b) they were likely originated by precipitation following low-temperature hydrothermal activity and submarine volcanism in an extensional geodynamic condition (i.e., an intra-back-arc basin setting) (El-Shazly et al., 2019) during Mozambique Ocean closure stage of the Pan-African Orogeny (∼696 Ma) (Abd El-Rahman et al., 2019) (c) other authors attribute their formation to the concomitant melting of glacial ice during the Snowball Earth’s interglacial period (∼750 Ma) (Stern et al., 2013), which is supported by the existence of small diamictite beds within the volcaniclastic successions (i.e., Wadi Kareim, Wadi Mobark, and Nuwaybah formations) (El-Shazly and Khalil, 2016). The majority of Egyptian BIFs are poorly constrained in terms of extent, grade, and tonnage, and in terms of tectonic, metamorphic, geochemical, and geochronological characteristics.

2.2 Deposit geology

The study area, Wadi El-Mishama, is located in the middle segment of the Eastern Desert (south of the Wadi El-Dabbah area, Figure 2). Wadi El-Mishama belongs to the Lower El-Dabbah Formation (total thickness of ∼2 km) that consists mainly of a metavolcano-sedimentary sequence and is overlain by subaerial sedimentary rocks (the Hammamat Group and Atshan Formation; Figure 3A). Three stratigraphic units are identified within the study area (i.e., arc-related volcanic flows, BIF-rich volcaniclastic rocks, and tuff-rich volcaniclastic rocks). The low-grade island-arc assemblages are commonly represented by massive pillow basaltic flows interlayered with a thin volcaniclastic sequence. In the upper part of this formation, pale greenish volcanic tuff along with several thin beds of well-preserved iron sequences (several meters thick) are observed. This succession contains well-preserved volcano-sedimentary structures including pillow lava, graded and parallel laminated tuffs, and iron oxide layers (Figure 3). Furthermore, the tuff-rich volcaniclastic layers display slumping, folding, and faulting features, indicating that these strata experienced syn-to post-depositional deformation.

2.3 Field observations and petrography

The BIF occurrences within the study area are notably significant, presenting several meters in thickness, which is indicative of extensive iron deposition events within the volcaniclastic rocks (Figure 3B). These BIF layers are characterized by their well-defined banding, alternating between iron-rich layers and silica-rich bands, reflecting periodic conditions of deposition (Figure 3C). The BIF outcrops exhibit varying sizes, with some extending several tens of meters in length and width and grade of these BIFs varies, with iron content significantly influencing the economic value of these deposits.

Microscopic analysis shows that the studied felsic lavas are mainly composed of feldspar and quartz. Euhedral magnetite is observed as fine-grained disseminations within these rock units (Figure 4). The well-exposed BIF (magnetite-rich beds alternating with jaspilites or silicate laminae) from the Wadi El-Mishama area is an example of Neoproterozoic oxide-facies BIFs in the ANS. They are commonly found at different scales, ranging from a few centimeters to as much as several meters thick. At the microscopic scale, the BIF layers are composed of alternating quartz-rich light and magnetite-rich dark microbands (Figures 4A, B). Magnetite and hematite are the common Fe-minerals in the studied BIFs (Figures 4C, D). Additionally, there are clearly developed and preserved sedimentary structures (i.e., bedding and graded bedding). The studied units are bounded by a strongly foliated metasedimentary sequence and are in intrusive contact with granitoids (hornblende–biotite granite) (Kiyokawa et al., 2020). The entire sequence experienced greenschist-facies metamorphism with brittle deformation and minor brittle-ductile deformation (i.e., pressure-temperature conditions of 373°C ± 61°C and 1.1–2.2 kbar) (El-Shazly and Khalil, 2016).

FIGURE 4

3 Investigative methodologies

3.1 Samples collection and preparation

A total of sixty (60) representative samples from BIFs and their host rocks distributed over the Wadi El-Mishama area were collected during two field trips. The location of the samples was highly dependent on the availability and accessibility of outcrops. The collected samples are massive, homogeneous, and least weathered, and were carefully selected based on color, texture, and macro-scale mineralogy variations. Twenty-five (25) samples were cut and polished to prepare thin and polished sections for carrying out systematic microscopic descriptions. Based on the detailed microscopic investigation, thirteen (13) samples were selected for whole rock chemical analysis, including eight (8) samples from BIFs and five (5) samples from metavolcanic wall rocks. These samples were also chosen for further analyses, that is, lead isotope analysis.

3.2 Analytical methods

Analyses performed at ALS Geochemistry, North Vancouver, British Columbia, Canada included the following: 1) major elements were analyzed by ICP-AES following lithium metaborate fusion; 2) trace elements (Ba, Cr, Cs, Ga, Hf, Nb, Rb, Sr, Ta, Th, U, V, Y, Zr, and the REE) by ICP-MS following lithium metaborate fusion (method ME-MS81); and 3) trace and some major elements (Ag, Al, As, Ba, Be, Bi, Ca, Cd, Ce, Co, Cr, Cs, Cu, Fe, Ga, Ge, Hf, In, K, La, Li, Mg, Mn, Mo, Na, Nb, Ni, P, Pb, Rb, Re, S, Sb, Sc, Se, Sn, Sr, Ta, Te, Ti, Tl, U, V, W, Y, Zn, Zr) by an ultra-trace four-acid digestion (HF, HClO₄, HCl, HNO₃) method (method ME-MS61L) followed by a mixture of ICP-AES and ICP-MS analysis. Data quality was monitored through the use of the following certified reference materials: OREAS-101b, 920, 47, 45e, and 14p; SY-4; BCS-512; AMIS0167, 0085, 0304. The REE anomalies of BIF samples were calculated as: Ce/Ce* = [(Ce_PAAS/(La_PAAS × Pr_PAAS)1/2]; Pr/Pr* = [Pr_PAAS/(Ce_PAAS × Nd_PAAS)1/2]; Eu/Eu* = [Eu_PAAS/(Sm_PAAS × Gd_PAAS)1/2]; La/La* = [(La_PAAS)/(3 × Pr_PAAS) –(2 × Ho_PAAS)] (Bolhar et al., 2004).

Lead isotopes were measured on a quadrupole ICP-MS at ALS Geochemistry following four acid digestion of whole rock powders (method MS61L-PbIS). Procedural blanks for most elements are less than the detection limits. For the Pb isotopes, procedural blanks are typically <0.01 ppm. Precision and accuracy of the Pb isotope results are typically 0.5%–3%, and for major and trace elements typically <5%.

For the major and trace elements a variety of certified reference materials (CRMs) were used including SY-4, BCS-512, AMIS0571, CDN-W-4, OREAS 920, OREAS-101b, AMIS0547, AMIS0167, AMIS0085, AMIS0304, OREAS 146, OREAS-14P, AMIS0461, OREAS 47, OREAS-45e. For the Pb isotopes, certified reference materials OREAS 920, 47, and 45e were analyzed concurrently. All CRM data is available on request.

4 Results

4.1 Whole rock geochemistry

4.1.1 BIFs

Major oxides, trace, and rare earth elements (REEs) of the analyzed samples from the Wadi El-Mishama iron deposits, combined with relevant geochemical parameters, are presented in Table 1. Whole-rock chemical analyses of BIFs compiled from previous studies of the ANS (Khalil et al., 2015; Abd El-Rahman et al., 2019; El-Shazly et al., 2019) (Supplementary Table S2), are used to discuss the depositional environment and origin of the studied BIFs.

The Fe₂O₃^t contents in the BIF samples range between 50.1 and 74.1 wt%, whereas the SiO₂ and CaO values vary from 26.6 to 34.4 wt% and from 1.3 to 7.7 wt%, respectively. The efficient removal of SiO₂, CaO, and MgO is responsible for the marked enrichment of iron in all ore samples (Gutzmer et al., 2008). Notably, all analyzed ore samples have significantly low concentrations of TiO₂ (0.08–0.2 wt%), Al₂O₃ (1.4–3.7 wt%), Na₂O (0.06–0.6 wt%), K₂O (0.07–0.3 wt%), MnO (0.03–0.05 wt%), and P₂O₅ (0.12–0.21 wt%, Table 1). On Harker diagrams (Figures 5, 6), the Fe₂O₃^t correlates negatively with SiO₂, CaO, Al₂O₃, P₂O₅, Cr, and Zr. There is no significant difference in major element contents between the Wadi El-Mishama BIFs and the Eastern Desert Algoma-type BIFs in the ANS (Figures 5, 6). The BIF samples reveal wide variations of Mn (182–336 ppm), Cr (20.5–118.5 ppm), Ni (9.8–61.3 ppm), and Sr (28.6–92.4 ppm) concentrations (Table 1). Furthermore, the contents of Zr (16–53 ppm), Hf (0.5–1.6 ppm), Cu (3.27–16.05 ppm), and Zn (7.8–25 ppm), are low in general. During hydrothermal, diagenetic, and weathering processes, aluminum and titanium are thought to be largely immobile (MacLean and Kranidiotis, 1987). Here, Al₂O₃ contents show an inverse relation with Fe₂O₃^T contents (Figure 5). In addition, during diagenesis, alumina-poor iron silicate minerals break down into iron hydroxide minerals and amorphous silica; as a result, chert bands are commonly formed where silica is trapped by iron oxide laminae. The SiO₂- Fe₂O₃^t biplot (Figure 5) shows a declining trend of SiO₂ with rising Fe₂O₃ values, displaying the presence of a hematite-rich component with small chert micro-bandings in some samples.

FIGURE 5

FIGURE 6

In the primitive mantle (PM) normalized trace element diagram, BIFs exhibit negative Ti, Sr, Nb, and Pr anomalies (Figure 7A). The total REEs (ΣREE) concentration varies between 56.00 and 75.85 ppm (Table 1). The studied samples are characterized by depletion of light REE relative to heavy REE (La/Yb_PAAS = 0.08–0.12) (Figure 7B). The studied BIFs show obviously weak positive Eu_PAAS anomalies (Eu/Eu* = 0.88–1.11, except for two samples with values of 0.67 and 0.73) and weak negative or no Ce_PAAS anomalies (Ce/Ce* = 0.83–0.93) (Table 1; Figure 7B). In addition, most of the BIF samples have strong positive La_PAAS anomalies (La/La* = 1.15–8.57). Compared to a chondritic Y/Ho value, the studied deposit has relatively high Y/Ho values (29.55–38.66, Table 1).

FIGURE 7

4.1.2 Metavolcanic rocks

The analytical results of the studied volcanic rocks are presented in Table 1. The bulk chemical compositions of hosting units for Algoma-type BIFs from the literature (i.e., Eastern Desert; Khalil et al., 2015; El-Shazly et al., 2019) are also used to discuss the possible source of parent magma and the formation setting of wall units (Supplementary Table S2).

4.2 Pb isotopic composition

Lead isotopic compositions for seven iron ore samples from the Wadi El-Mishama deposit are presented in Supplementary Table S3. ²⁰⁶Pb/²⁰⁴Pb values range from 16.67 to 19.80 with an average of 17.71. ²⁰⁷Pb/²⁰⁴Pb values are between 14.33 and 16.00 with an average of 15.38; whereas ²⁰⁸Pb/²⁰⁴Pb values are from 34.67 to 39.20 with an average of 37.27.

5 Discussion

5.1 Depositional environment

The depositional environment of BIFs has been a subject of considerable debate. Regionally, it has been proposed that the most likely tectonic environments from which iron may have been released and then accumulated as BIFs are mid-oceanic ridges or hot spots (Dalstra et al., 2008). Algoma-type BIFs are most commonly found in Precambrian successions that contain a significant proportion of submarine volcanic rocks. These stratigraphic horizons were formed in oceanic arc and back-arc environments (Peter, 2003). More especially, numerous giant and world-class BIFs were created during intra-plate tectonic-thermal and rifting episodes (Pirajno and Bagas, 2008). It is well known that the Precambrian Algoma-type BIFs are stratigraphically and genetically linked to VMS mineralization in many cases worldwide (e.g., Qingyuan greenstone belt, North China and Abitibi greenstone belt, Canada) (Taner and Chemam, 2015; Peng et al., 2022). The precipitation of iron ores may be found close to or distal from the high-temperature up-flow zone of VMS-related submarine magmatic-hydrothermal activity (Peter et al., 2014). Locally, the iron successions in the Egyptian Eastern Desert have been linked to volcanic eruptions and are proposed to originate in an intra-back-arc basin setting (Faisal et al., 2022).

Based on the Nb/Y versus Zr/Ti classification diagram (Figure 8A) (Winchester and Floyd, 1977), the studied host rocks showed variance between andesitic to rhyolitic compositions. They have a transitional affinity (tholeiitic - calc-alkaline) on the Zr/Y vs. Th/Yb diagram (Figure 8B). The tectonic discrimination diagrams show clearly that these rocks mainly plot within island arc setting fields, indicating the oceanic nature of the BIF host rocks (Figures 8C, D). The high field strength elements (HFSEs, e.g., Zr, Hf, Nb, Ta, P, and Ti) and HREE values are helpful markers for determining potential tectonic settings of extrusive rocks (Pearce, 2008). In the Ti against Zr tectonic discrimination diagram (Figure 8C), the host rocks of Wadi El-Mishama BIFs lie mainly in the island arc field. According to the Th/Yb versus Nb/Yb diagram (Figure 8D), the studied units have an affinity to oceanic island arcs. This suggests that the iron sequence of the Wadi El-Mishama area is typical of an intra-oceanic arc setting (i.e., back-arc basin). This result is in agreement with the depositional environment of island arcs from the Egyptian Eastern Desert which, were deposited in a back-arc spreading tectonic setting during the closure of the Mozambique Ocean (Faisal et al., 2020).

FIGURE 8

5.2 Source characteristic and depositional process

5.2.1 Detrital input

Iron formations are typically detritus-free, however, terrigenous contamination has been documented throughout the world (Hou et al., 2019). In general, some oxides (Al₂O₃ and TiO₂), transition metals, HFSEs, and LILE are suitable monitors for detecting terrigenous contamination. During hydrothermal, diagenetic, and weathering processes, aluminum and titanium are thought to be largely immobile (MacLean and Kranidiotis, 1987). The studied BIFs exhibit relatively low Al₂O₃ + TiO₂ contents ranging from 1.46 to 3.81 wt.% (Table 1), with a positive correlation between them (Figure 9A) suggesting the presence of terrigenous components during their deposition (Barrote et al., 2017). In addition, a clastic component of igneous origin may have been present in the examined samples due to relatively high Cr (20.5–118.5 ppm, Table 1) and Zr (16–53 ppm) values (Bolhar et al., 2004).

FIGURE 9

The Y/Ho (29.55–38.66) values of the studied BIFs, display little admixture of clastic and volcanic components during their deposition (Bolhar et al., 2004). The BIF samples have Eu anomalies of 0.88–1.11 (Figure 7B), consistent with low terrigenous input. Relatively higher values of Sr and Y (Table 1) indicate the presence of a source of crustal felsic rocks (i.e., these components are generally produced from the weathering of crustal felsic rocks) (Rao and Naqvi, 1995). Thus, all these features suggest that the Wadi El-Mishama iron deposit was variably influenced by detrital materials. The studied BIFs samples plot close to the East Pacific Rise hydrothermal metalliferous sediments field in the Fe/Ti versus Al/(Al+Fe+Mn) diagram (Figure 9B), which suggests that they are characterized by a low proportion (less than 20%) of detrital components. It is consistent with low U/Th values (i.e., Wadi El-Mishama BIFs vary from 0.15 to 0.37) that are typical of relatively pure chemical iron formations (Thurston et al., 2012).

5.2.2 Hydrothermal activity

The immobile elements play a key role in understanding metallogenetic evolution, and the REEs signature of iron deposits has been employed as a proxy to determine the origin of silica and iron (Soh Tamehe et al., 2018). The LREE depletion, HREE enrichment, and positive La anomalies of the studied BIFs indicate seawater signatures to some extent (Bolhar et al., 2004; Soh Tamehe et al., 2018). U enrichment, which is normally enriched in the evolved crust, is usually reflected a detrital contribution to BIFs deposits (Taylor and McLennan, 1985; Sylvestre et al., 2017). In contrast, U depletion is mostly linked to oxic, early-diagenetic alteration resulting in the release of U back into seawater (Dunk and Mills, 2006). The BIFs samples contain relatively lower concentrations of U (0.06–0.15 ppm) compared to Th (0.27–0.94 ppm), which is indicative of their exposure to seawater with a slow rate of accumulation (Roy and Venkatesh, 2009). On the REE patterns normalized PAAS plot (Figure 7B), the samples lie between the average composition of low-temperature hydrothermal solutions and South Pacific seawater. The Zr/Hf against Y/Ho plot (Figure 10A) signifies that most of the data are super-chondritic (Y/Ho > 27.7), pointing to the interaction of hydrothermal fluids with seawater as well as certain potential sources of hydrogenous-detrital contamination (i.e., volcanic input or terrigenous clastic materials from country rocks). The ratios of Eu/Sm and Sm/Yb reflect the rates of reaction between hydrothermal fluids (0.1%–5%) and seawater during the deposition of BIFs (Figure 10B; Alexander et al., 2008), that is, a 1000:1 to 100:5 mixture of seawater and high-temperature hydrothermal fluids. The ratio of seawater to hydrothermal fluids is further emphasized by using the discrimination diagrams of Alexander et al. (2008) (e.g., Sm/Yb vs. Eu/Sm; Figure 10B), where samples accumulate close to seawater and low-temperature hydrothermal fluids (<0.1%) fields and far away from high-temperature hydrothermal fluids.

FIGURE 10

The obvious positive Eu anomalies in Archean and Paleoproterozoic BIFs suggest that sub-oceanic waters were affected by hydrothermal fluids from deep-sea spreading sites (e.g., Bau and Dulski, 1996; Bolhar et al., 2004; Planavsky et al., 2010). More specifically, high-temperature hydrothermal fluids (>250°C) are characterized by strong positive Eu anomalies, whereas low-temperature fluids (<200°C) are characterized by weak to no positive Eu anomalies (Basta et al., 2011). Variations in the hydrothermal fluid ratios to seawater, as well as the mobilization of Eu at temperatures above 200°C during the alteration or metamorphism of the ores, could have caused changes in the intensity of the Eu anomalies (Bolhar et al., 2004). So, the weak positive Eu anomalies of the studied BIFs are consistent with low-temperature (<250°C) shallow hydrothermal volcanic exhalative systems.

The high Y/Ho values in some fluids can result from the mixing of seawater and hydrothermal fluids during REE scavenging by iron (oxy) hydroxide minerals. As mentioned previously, the Y/Ho values in the studied BIFs are super-chondritic (Y/Ho > 27.7), and similar to those of most Neoproterozoic BIFs (mean Y/Ho = 29.2; n = 77) (Cox et al., 2013). Additionally, this does not imply that a high-temperature hydrothermal input occurred. The iron samples from the study area have low Cu (3.27–16.05 ppm) and Zn (7.8–25 ppm) concentrations that are relatively near to the crustal background levels of these metals (Angerer et al., 2012), indicating the studied BIFs deposit was likely formed from low-temperature (<250°C) hydrothermal fluids without significant accumulation of Cu-Zn massive sulfide or were formed distal from the high-temperature up-flow zone of VMS-related magmatic-hydrothermal system.

The Fe/Ti versus Al/(Al + Fe + Mn) diagram reveals the ratios of hydrothermal fluids to the detrital contamination of clastic inputs during BIF deposition. All collected samples plot close to the field of East Pacific Rise hydrothermal sediments (Figure 9B). Due to the low Al₂O₃ content (1.38–3.65) of the examined iron ores, hydrothermal solutions rather than detrital mineral assemblages were responsible for the weak positive Eu_PAAS anomalies. Also, these Eu_PAAS anomalies signify that the fluids that carried SiO₂ and FeO to the ambient ocean were derived from a reducing environment (Campbell et al., 1988). Furthermore, the hosting volcanic lavas display positive Eu anomalies (Table 1).

In the Co/Zn versus Fe₂O₃^T diagram (Figure 10C), all samples were plotted on the hydrothermal field (Toth, 1980). Klein and Beukes (1989) developed the relationship between ΣREE and Σ (Co + Cu + Ni) of hydrogenous and hydrothermal deposits to construct the fields of these deposits (Figure 10D). According to this diagram, the majority of the studied BIFs are located close to the area where hydrothermal deposits are present. This suggests that a significant portion of the iron in the investigated BIFs was added to the seawater by hydrothermal solutions derived from hydrothermally active marine environments. All BIF samples are shown to plot on the field of hydrothermal origin in the Fe–Al–Mn triangular discrimination diagram (Figure 11A). Therefore, we propose that the Wadi El-Mishama iron deposit was precipitated from a mixture of seawater and low-temperature hydrothermal fluids (<200°C) with variable detrital contamination.

FIGURE 11

5.2.3 Redox conditions

It has been suggested that the ancient oceans at the beginning of the Neoproterozoic Era were dominantly in an anoxic state (Tostevin et al., 2016). The Ni/Co and V/(V + Ni) values are commonly used to evaluate the paleo-redox conditions (Jones and Manning, 1994). It has been shown that Ni/Co values of <5 indicate oxic conditions, whereas the values of >5 exhibit dysoxic states (i.e., very low oxygen concentration, between anoxic and hypoxic) in the deposition area (Jones and Manning, 1994). Ni/Co values of the analyzed iron samples display dysoxic conditions (avg.= 6.11). On the other hand, V/(V + Ni) values from 0.50 to 0.60 suggest oxic conditions, whereas values from 0.60 to 0.85 represent anoxic conditions (Rimmer, 2004). In the present study, V/(V + Ni) values have an average of 0.59, which is at the interface between oxic and anoxic conditions. Ce responds to changes in the redox state. Compared to shale composites, oxic water has a negative Ce anomaly, whereas the underlying anoxic water lacks this negative anomaly and may exhibit a small positive Ce anomaly (de Baar et al., 1988). The collected ores display weak negative or no Ce_PAAS anomalies, implying suboxic to oxic conditions. A lack of a noticeable Ce anomaly in hydrothermal iron-rich crusts (modern basins) was attributed by Stoffers et al. (1993) to the mixing of hydrothermal fluids with seawater. Furthermore, Ce/Ce* values below ∼0.10 suggest oxic conditions, whereas the ratios above 0.10 reflect anoxic conditions (Wright et al., 1987). The data from the study area have Ce/Ce* ranging from 0.83 to 0.93 which clearly shows anoxic events during the formation of iron deposits. Moreover, it has been proposed that a negative Ce anomaly can be divided into three groups: (a) ∼0.9–1.0; (b) ∼0.6–0.9; and (c) <0.5 which represents anoxic, suboxic, and oxic marine water, respectively (Chen et al., 2015). The studied BIF samples plot mainly on the positive La anomaly field (with no distinguished Ce anomaly) in the (Pr/Pr*)_PAAS vs. (Ce/Ce*)_PAAS diagram (Figure 11B) and completely in the field of late Paleoproterozoic BIFs deposits, which reflect a suboxic to anoxic condition during BIF precipitation (Wang et al., 2016). Conclusively, the Ce/Ce* values documented by the investigated BIFs reflect that the sediments were mainly deposited at the transition between suboxic and anoxic conditions. The U/Th values of >1.25 signify anoxic conditions (Nath et al., 1997). The U/Th values of the studied BIFs vary from 0.15 to 0.37, which are consistent with suboxic to oxic conditions. In addition, on the PAAS-normalized diagram, oxic water is commonly characterized by HREE enrichment compared with anoxic water (German et al., 1991). LREE are more susceptible to removal from solution through adsorption reactions, while HREE are almost entirely bound by stable carbonate complexes (Byrne and Kim, 1990; Byrne and Sholkovitz, 1996; Alibo and Nozaki, 1999). For example, cerium displays a strong negative anomaly On the PAAS-normalized REE pattern for an oxygenated marine environment (Ce^III oxidized to Ce^IV), while no noticeable negative anomaly in a suboxic and anoxic environment (Byrne and Sholkovitz, 1996; Ohta and Kawabe, 2001). LREE-depletion relative to HREE in the PAAS-normalized diagram of the studied BIFs samples (Figure 7B) suggests the deposition has occurred under slightly oxidized seawater. Furthermore, several metals are commonly to be complexed by natural organic substances, for example, zinc is largely complexed in the upper water column than copper (Bruland, 1989; Donat and Bruland, 1990). Thus, low Cu/Zn values largely signify oxidizing environments in the depositional basin, whereas high Cu/Zn values imply reducing conditions (Hallberg, 1976; Nagarajan et al., 2007; Ngueutchoua et al., 2017). The Cu/Zn values of the studied ores range from 0.24 to 1.37, which are suggestive of oxic conditions. Consequently, the above-mentioned geochemical signatures indicate that the sedimentary environment of the Wadi El-Mishama BIFs was primarily dysoxic to oxic.

5.3 Comparison with other Algoma-type BIFs

A comparative study of the major and trace element contents of Wadi El-Mishama BIFs with BIFs from other regions within the ANS (Figures 5, 6; Supplementary Table S2) shows that there is a significant degree of consistency and similarity. The SiO₂–Al₂O₃–Fe₂O₃^T (Figure 11C) (Govett, 1966) ternary diagrams show that the studied iron ores plot closer to the Fe₂O₃ apex and within the field of Algoma-type Precambrian BIFs, thus implying a closer chemical similarity to other Precambrian iron deposits. Some iron samples from BIFs of the Gebel Semna, Um Ghamis, and Wadi Kareim areas are variable in their REE patterns due to weathering and metamorphism effects. No notable positive Eu anomalies were found in the examined iron ores (Eu/Eu_PAAS = 0.88–1.11; Figure 7B), which is consistent with other Neoproterozoic iron deposits. It is noteworthy that Algoma-type iron deposits typically exhibit a weak negative Ce anomaly that resembles seawater (Aftabi et al., 2021). The investigated iron formations primarily exhibit negligible to small negative Ce anomalies (Ce/Ce_PAAS = 0.83–0.93) that are similar to those recorded in the ANS iron deposits (Figure 7B).

Genetically, two hypotheses have been proposed for the formation of ANS BIF deposits. The first view is that they have been linked to global glaciations (Snowball Earth episodes) similar to Rapitan-type BIF. More specifically, Stern et al. (2006) and Ali et al. (2010) attributed the formation of iron deposits at Wadi Kareim, Central Eastern Desert (Figure 1, location no. 6) to the Sturtian glaciation event with evidence of glacial diamictite deposited with iron ores. The lack of glacial/ice-transported detritus (sedimentological evidence of glaciation) in the study area suggests that the Snowball Earth hypothesis is not fully convincing for the investigated iron formations. The other model is that they were formed in a tectonically and magmatically active basin during the closure of the Mozambique Ocean (Abd El-Rahman et al., 2019; El-Shazly et al., 2019; Kiyokawa et al., 2020). The extensive association of the volcanic and volcaniclastic rocks with the iron formations is a common observation in this model.

The Wadi El-Mishama BIFs are intimately associated with volcanic units that formed within an intra-oceanic arc setting based on the resultant tectonic discrimination diagrams (Figures 8C, D). As a result, the iron was most likely sourced from hydrothermal vent fluids. Moreover, these features are largely in agreement with those proposed in previous studies in the Egyptian Eastern Desert (Abd El-Rahman et al., 2019; El-Shazly et al., 2019). This conclusion is also supported by relatively homogeneous, low radiogenic Pb isotopic ratios from the study area which plotted close to the field of BIFs from the ANS and within the oceanic field on the ²⁰⁶Pb/²⁰⁴Pb versus ²⁰⁸Pb/²⁰⁴Pb binary diagram (Figure 12A). This is a sign that the Pb metal was derived from geological bodies that had homogeneous Pb isotopes or those that had varied Pb isotopes but were homogenized prior to precipitation within extensional geodynamic regimes (i.e., an intra-oceanic arc environment) over a relatively restricted period. Besides, the current isotopic data lies in the area between the mantle and the crust (Figure 12B) and overlaps that of the Pb isotopic composition of MORB (Figure 12C). It demonstrates that Pb metal may have originated from the same source of iron and base metals (i.e., magmatic-hydrothermal fluids and leaching of basement lithologies).

FIGURE 12

Based on the above discussion, we suggest that Wadi El-Mishama BIFs have geochemical signatures of modern iron-rich crusts that their formation is connected to low-temperature shallow hydrothermal volcanic exhalative systems related to suboxic to oxic basins. Lastly, the Wadi El-Mishama and other BIFs from the ANS have quite similar geochemical fingerprints, which could point to a similar mechanism of formation from hydrothermal vent fluids.

6 Concluding remarks

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