Ironstone deposits hosted in Eocene carbonates from Bahariya (Egypt)—New perspective on cherty ironstone occurrences

    Ironstone deposits hosted in Eocene carbonates from Bahariya (Egypt) – new perspective on cherty ironstone occurrences A.M. Afify, M.E. Sanz-Montero, J.P. Calvo PII: DOI: Reference:

S0037-0738(15)00203-1 doi: 10.1016/j.sedgeo.2015.09.010 SEDGEO 4911

To appear in:

Sedimentary Geology

Received date: Revised date: Accepted date:

3 August 2015 17 September 2015 18 September 2015

Please cite this article as: Afify, A.M., Sanz-Montero, M.E., Calvo, J.P., Ironstone deposits hosted in Eocene carbonates from Bahariya (Egypt) – new perspective on cherty ironstone occurrences, Sedimentary Geology (2015), doi: 10.1016/j.sedgeo.2015.09.010

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Ironstone deposits hosted in Eocene carbonates from Bahariya (Egypt) – new perspective on cherty ironstone occurrences

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Afify, A.M.a,b,*, Sanz-Montero, M.E.a, Calvo, J.P.a

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a) Department of Petrology and Geochemistry, Faculty of Geological Sciences, Complutense

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University, Madrid, C/ José Antonio Nováis, 2, 28040 Madrid, Spain

b) Department of Geology, Faculty of Science, Benha University, 13518 Benha, Egypt

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* Corresponding author ([email protected])

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ABSTRACT

This paper gives new insight into the genesis of cherty ironstone deposits. The

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research was centered on well-exposed, unique cherty ironstone mineralization

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associated with Eocene carbonates from the northern part of the Bahariya Depression (Egypt). The economically important ironstones occur in the Naqb Formation (Early

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Eocene), which is mainly formed of shallow marine carbonate deposits. Periods of lowstand sea-level caused extensive early dissolution (karstification) of the

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depositional carbonates and dolomitization associated with mixing zones of fresh and marine pore-water. In faulted areas, the Eocene carbonate deposits were transformed into cherty ironstone with preservation of the precursor carbonate sedimentary features, i.e. skeletal and non-skeletal grain types, thickness, bedding, lateral and vertical sequential arrangement, and karst profiles. The ore deposits are composed of iron oxyhydroxides, mainly hematite and goethite, chert in the form of micro- to macro-quartz and chalcedony, various manganese minerals, barite, and a number of subordinate sulfate and clay minerals. Detailed petrographic analysis shows that quartz and iron oxides were coetaneous and selectively replaced carbonates, the

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coarse dolomite crystals having been preferentially transformed into quartz whereas the micro-crystalline carbonates were replaced by the iron oxyhydroxides.

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A number of petrographic, sedimentological and structural features including

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the presence of hydrothermal-mediated minerals (e.g., jacobsite), the geochemistry of

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the ore minerals as well as the structure-controlled location of the mineralization suggest a hydrothermal source for the ore-bearing fluids circulating through major faults and reflect their proximity to centers of magmatism. The proposed formation

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model can contribute to better understanding of the genetic mechanisms of formation

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of banded iron formations (BIFs) that were abundant during the Precambrian.

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Keywords:

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carbonates, Egypt.

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Cherty ironstone, dolomitization, tectonic constraints, hydrothermalism, Eocene

1. Introduction

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The Eocene strata in northern Bahariya contain ironstone deposits of economic significance, some of them reaching a large size (Fig. 1). Despite significant research on these deposits (El Shazly, 1962; El Akkad and Issawi, 1963; Basta and Amer, 1969, and references therein), their origin is still a matter of debate. More recent publications show different and contrasting hypotheses for the source and mechanisms of formation of the Bahariya ironstones. Dabous (2002) concluded that the Bahariya ironstone deposits are not lateritic and that their formation was related to mixing of warm ascending groundwater leaching iron from the underlying Nubia aquifers and descending water with iron leached from the overlying Upper EoceneLower Oligocene glauconitic clays. Salama et al. (2013, 2014) concluded that the

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Bahariya ironstones were deposited primarily in a marine setting and their formation was enhanced by microbial activity. They related the formation of the ironstone to

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global warming during the early Paleogene, closely associated with eustatic sea-level

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changes. Recent work by Baioumy et al. (2013, 2014) supports sources and

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mechanisms of iron and manganese formations that are contradictory: supergenetic ore deposits (most probably from the Naqb limestone host rock) or hydrogenous iron mixed with iron of hydrothermal origin (sea water precipitation to hydrothermal

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exhalite).

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New insight based on field, petrographic, mineralogical and geochemical studies of the ironstone deposits in the northern part of the Bahariya Depression (Fig.

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1) is provided in this paper. The similarities between the ironstones and carbonate

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host rocks as well as the close relationship between the ore mineral body and the regional tectonic structure led us to revisit the models proposed for the formation of

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ironstone deposits of Bahariya. Moreover, the scarcity of Phanerozoic, in particular Cenozoic cherty ironstone makes the Bahariya ore deposits an interesting case study

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to investigate some new perspectives about the formation constraints of these kinds of rocks. It can help also in understanding the mechanisms of older iron-rich deposits where chert is an important constituent. In the study area, iron is paired with quartz formation, this resulting in sedimentary structures that resemble some of those present in banded iron formations (BIFs). The term cherty ironstone is referred here to the richness of quartz, usually higher than 30% in the ore deposits.

2. Geologic setting The Bahariya Depression is located near the central part of the Western Desert of Egypt (Fig. 1). The depression shows an elliptical geometry surrounded by a

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carbonate plateau. The stratigraphic succession exposed in the northern part of Bahariya comprises the Bahariya Formation (Early Cenomanian), El Heiz Formation

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(Late Cenomanian) and Hefuf Formation (Campanian) forming the floor of the

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depression and surrounded by a carbonate plateau of Eocene rocks (El Akkad and

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Issawi, 1963, Said and Issawi, 1964). The Eocene carbonate rocks start with the Naqb Formation, which overlies unconformably the siliciclastic deposits of the Upper Cretaceous Bahariya Formation (Figs. 1, 2; Said, 1962; Afify et al., 2015a). The Naqb

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Formation is overlain by the Qazzun and El Hamra formations, which are poorly

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represented with reduced thickness at the northeastern part of the study area (Figs. 1, 2). The ironstone deposits occur associated with the Eocene carbonates in three areas

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(Fig. 1). Both the Eocene units and the ironstone deposits are unconformably overlain

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by the Oligocene ferruginous quartzarenite beds of the Radwan Formation (Figs. 1, 2). Outcrops of basaltic to doleritic igneous dykes, sills, laccoliths and lava flows of

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middle Miocene age (El-Etr and Moustafa, 1978; Meneisy, 1990) can be observed in the northern part of the depression and to the south of the study area. The Bahariya

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basalts are of alkaline type. Two varieties, i.e. olivine basalt and dolerite were distinguished (Meneisy, 1990). The Bahariya Depression is punctuated by a NE-trending right-lateral wrench

fault system (Fig. 1), which is associated with several doubly plunging folds and extensional faults (Sehim, 1993; Moustafa et al., 2003). Three phases of structural deformation affected the northern part of the depression: (1) post Campanian – preMiddle Eocene ENE right-lateral transpression, (2) post-Eocene reactivation and (3) Middle Miocene extensional deformation (Moustafa et al., 2003). The strain regime in the Bahariya area was transpressional and led to the formation of the Bahariya swell by the combined effect of the first and second

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deformation events (Said and Issawi, 1964). The stresses that formed the NE-SW doubly plunging anticline folds and ENE strike-slip faults characteristic of the first

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deformation phase continued throughout the Paleocene and Eocene. A similar

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evolution has been documented in the Esna Formation (Paleocene-Early Eocene;

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Sanz-Montero et al., 2013; El Ayyat, 2013) at the northern part of the Farafra Depression. Moreover, syndepositional tectonic activity and seismic pulses took place

and Issawi, 1964; Obaidalla et al., 2006).

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during the deposition of the Eocene sediments in the Bahariya and Farafra areas (Said

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The second tectonic phase occurred after the Middle Eocene and before the Oligocene. Tectonic inversion continued in the area, leading to the development of

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folds and small domes in the Eocene formations (Said and Issawi, 1964; Moustafa et

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al., 2003). As a consequence, the carbonate sequence was fractured and folded along NE to ENE oriented right-stepped en-échelon folds (Fig. 1). The faulting pattern

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shows major NE-SW dextral strike-slip faults and local thrusts associated with the folds, e.g., southern part of Ghorabi area and at El Harra area (Fig. 1), and WNW left-

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stepped, en-échelon normal faults and E-W normal faults. Some of the latter faults affected the Oligocene Radwan Formation and were mostly related to the Middle Miocene extensional deformation. Associated volcanic activity is represented by erupted and/or intruded materials through fissures and discontinuities. This extensional phase was most probably related to the opening of the Gulf of Suez-Red Sea rift leading to the separation of Arabia from Africa (Moustafa et al., 2003). The second and third phases were the main tectonic events that affected the carbonates forming the plateau in the Bahariya region.

3. Materials and methods

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Field work was focused on detailed description and sampling of the stratigraphic succession forming the plateau at the northern part of the Bahariya area.

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About 140 samples of carbonate rocks, ironstones, sandstones, clays and other rocks

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were collected. Up to 120 indurated samples were prepared as thin sections and

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polished slabs. Petrographic characteristics were determined using an Olympus BX51 optical microscope with white light and ultraviolet fluorescent light sources as well as a Nikon reflected light microscope. The staining method of Lindholm and Finkelman

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(1972) with alizarin red-potassium ferricyanide was used to differentiate between

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carbonate minerals. For high-resolution textural and morphometric analyses, carboncoated thin sections and fresh broken pieces were studied using electron microprobe

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analysis (EMPA) and scanning electron microscopy (SEM). A subset of 21 carbon-

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coated thin sections were prepared for BSE (backscattered images), SE (secondary images) and elemental analyses (in wt. %) on a JEOL Superprobe JXA 8900-M

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wavelength dispersive electron microprobe analyzer (WDS-EMPA) equipped with four crystal spectrometers and beam diameter between 2 to 5 µm to minimize damage

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from the electron beam. SEM study was carried out on 41 fresh pieces placed on sample holders supported by carbon conductive tape, followed by sputter coating of gold and studied with a JEOL JSM-820 operating at 20 kV and equipped with an Oxford energy dispersive X-ray microanalyzer (SEM-EDAX). Mineral compositions for nearly all the collected samples were verified by XRD analyses that were performed using a Philips PW-1710 diffractometer under monochromatic Cu- Kα radiation (λ= 1.54060 Å) operating at 40kv and 30 mA, a step size of 2θ is 0.02º and time per step of 2 s. XRD analyses were performed following the method of Chung (1974) using EVA Bruker software. Eighteen samples were mechanically crushed for geochemical analyses using EDXRF (energy dispersive X-ray fluorescence). Fused

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discs were prepared for these samples and analyzed for their major oxides and trace elements using a Bruker S2 RANGER X-ray fluorescence spectrometer with X-Flash

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Silicon Drift Detector. Loss on ignition (L.O.I.) was obtained by heating 1g of

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powdered sample at 1000 °C for 1 h. Carbon and oxygen isotopic compositions for

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dolomite were reported for 15 samples using standard δ notation in units of ‰ relative to VPDB standard. δ13C and δ18O values were measured on CO2 released from differential dissolution of 10–20 mg of washed sample in 100% H3PO4. Calcite was

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after an additional 24 hours step at 70° C.

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removed as CO2 after 4 hours of reaction at 25° C. For dolomite, CO2 was extracted

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4.1. Carbonate host rocks

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4. Results

Outcrop observation of the Eocene formations in the northern part of the

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Bahariya Depression shows that the carbonate rock units were totally replaced and/or cemented by iron-bearing minerals and/or quartz in the vicinity of major faults (Table

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1; Figs. 1–3). Most descriptions and interpretations in this paper deal with the Naqb Formation as it constitutes the major rock unit hosting the iron ore bodies.

4.1.1. Sedimentology The carbonate deposits of the Naqb Formation were studied from 4 sections located near Ghorabi and El Harra (Fig. 1) measuring up to 13 m-thick (Fig. 3). The carbonate deposits are mostly dolostone, but the depositional fabrics are preserved well enough to allow description of the primary carbonate features, i.e. type of skeletal and non-skeletal particles, depositional fabrics, sedimentary structures and bedding geometry (Fig. 4A–E). The carbonate deposits of the Naqb Formation are

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fossiliferous showing abundant tests of nummulites, alveolinids, miliolids, gastropods, bivalves, dasycladacean algae and echinoids. Based on the foraminiferal

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content, the Naqb Formation is Late Ypresian in age (Boukhary et al., 2011).

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The Naqb Formation is subdivided into two sequences separated by a

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paleokarstic surface that can be used as a marker horizon in both the carbonates and their equivalent ironstone deposits (Figs. 2, 3, 4A). The paleokarst surface was described previously by Salama et al. (2014) who attributed the ironstone deposits

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underlying and overlying the surface to the Naqb and Qazzun formations,

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respectively. Observations at a regional scale show, however, that the two carbonate sequences separated by the paleokarst belong to the Naqb Formation. Brecciation as

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well as formation of speleothems with small- to medium-scale dolines, solution caves

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and sinkholes and irregular concentric laminae are common features in this unit (Fig. 4E–I). Facies description and interpretation of this rock unit are summarized in Table

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1. Three main facies can be distinguished in the lower sequence (Table 1; Fig. 3), i.e. nummulitic dolostone/marly dolostone (F1; Fig. 5A), thick-bedded fossiliferous (and

5E, F).

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oolitic) dolostone (F2; Fig. 5B–D) and massive non-fossiliferous dolostone (F3; Figs.

The upper sequence is composed of dolostone beds which locally show

pseudomorphs of evaporite minerals, likely sulfates. Three main facies can be distinguished (Table 1; Figs. 3, 6A–D), i.e. thick-bedded fossiliferous dolostone (F4), stromatolitic-like laminated dolostone (F5) and bivalve dolostone (F6). Sedimentary features, i.e. rosette-like and laminated evaporites, rhizoliths and dissolution tubes, laminated fenestral fabrics associated with desiccation cracks, dissolution, brecciation and calcretization characterize this sequence (Fig. 6A–C). Calcretes exhibit biogenic features such as rhizoliths and alveolar septal structures. Carbonate lamination,

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occurrence of bird´s eyes structures and vuggy pores are considered to represent stromatolite fabrics (Figs. 6B, C).

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Field and petrographic observations allowed recognition of superimposed

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karstic features pointing to development of two karstification phases in the Naqb

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Formation. The first karstification phase was related to subaerial exposure of the lower sequence, this resulting in the development of a variety of karstic features, e.g., pseudospherulitic fibrous, fan-like and broom-like carbonates that were later

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dolomitized (Figs. 6E, F). Karst features related to the second karstification phase are

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recognizable in the two carbonate sequences of the Naqb Formation as well as in the carbonates of the overlying Qazzun and El Hamra formations, where vertical, inclined

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and sub-horizontal dissolution tubes and fractures are either filled by quartz and/or

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large calcite crystals (Fig. 6F).

The paleoenvironmental interpretation of the carbonate facies of the Naqb

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Formation is sketched in Fig. 7. The thickening-upward sequence characteristic of the lower unit reflects a shallowing upward trend from facies F1 to F3, which is also

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marked by the dominance of nummulites and alveolinids at the bottom and an increase of textularids and miliolids towards the top (Beavington-Penney and Racey, 2004). The shallowing upward evolution of the lower sequence culminated with karst development after exposure. Facies F4, F5 of the upper sequence are dominated by thin laminated fabrics (stromatolites), desiccation cracks, rhizoliths, and scarcity of fossils representative of very shallow intertidal-supratidal conditions (Tucker and Wright, 1990; Suárez-González et al., 2015). The carbonate deposits forming the uppermost part of the upper sequence show bivalve packstone fabrics representative of intertidal-shallow subtidal environments with open marine conditions. Thus, the Eocene carbonate deposits record eustatic sea-level changes.

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The carbonate deposits of the Naqb Formation underwent extensive diagenetic processes, i.e. micritization, compaction, dissolution and karstification and

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dolomitization (Figs. 5, 6) which were followed by chertification, formation of iron

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oxyhydroxides, Mn-bearing minerals, and associated gangue minerals in areas where

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the carbonate host rocks were faulted. Timing, paragenesis and relationship of these processes are discussed below.

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4.1.2. Mineralogy and geochemistry of carbonate host rocks

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The main minerals determined in the carbonate deposits of the Naqb Formation are dolomite and quartz, with minor amount of calcite. The dolomite

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content ranges from 30% to more than 90% whereas the calcite content reaches

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locally up to 15%, especially in the non-silicified samples. Quartz content is higher where the carbonate deposits approach the iron mineralized beds.

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According to XRF analyses, the dolostones are mainly composed of CaO (mean: 24%), MgO (mean: 13%), Fe2O3 (mean: 2.2%), MnO (mean: 1.3%) and SiO2

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(mean: 34.88%; Table 2). High SiO2 and Fe2O3 contents were determined in ferruginous and silicified dolostone samples close to the mineralized areas. The probe analyses of individual dolomite crystals from the different facies (Table 3) show nearly stoichiometric non-ferroan dolomite and give the following values: CaO (mean: 30.23%), MgO (mean: 21.6%), FeO (mean: 0.65%), MnO (mean: 0.1%), and SrO (mean: 0.088%). The highest Fe2O3 content was determined from pore-filling hematite and it is not considered representative for iron in dolomites. Similarly, strontium content using probe analyses (0.002–0.173%; Table 3) is lower than that determined by XRF (0.069–2.264%; Table 2), which is probably influenced by porefilling Sr-rich minerals (e.g., recent evaporitic cements).

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The stable isotope compositions of the dolostone samples show a good covariant trend between δ13C and δ18O values referring to the VPDB standard (Table

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2; Fig. 8). Light δ13C (-5.46‰ to -1.05‰) and δ18O values (-7.71‰ to -2.74‰) were

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determined in carbonate samples from the lower sequence of the Naqb Formation

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whereas the isotope compositions are slightly heavier in carbonate samples from its upper sequence (-1.69‰ to +0.38‰ for δ13C and -4.33‰ to -0.77‰ for δ18O).

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4.2. Relationships between Eocene carbonate host rocks and ironstones

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Extensive replacement of the Eocene carbonate formations by Fe and Mn oxides and quartz occurs in localized areas near major fault lineaments where the

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carbonate rocks are altered partially to totally into cherty ironstone (Table 1; Fig. 3).

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Yet the cherty ironstone deposits retain many stratigraphic and sedimentary features of the Naqb carbonate rocks, i.e. thickness (from 7 to 13 m), bedding, and lateral and

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vertical sequential arrangement, thus allowing correlation between the ironstone and carbonate deposits and providing evidence for the replacement origin of the former

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(Figs. 3, 9A, B). The main lithostratigraphic features and facies of the host carbonate deposits are well-preserved where replaced by iron (e.g., the stromatolitic fabrics; Fig. 9C), with preservation of some unaltered thin clay/marl interbeds (Fig. 9D). With increasing distance from the major faults, the mineralization fades away, being localized only along the sedimentary discontinuities (Figs. 4A, B). Beyond these points, the carbonate deposits are not replaced by iron, even though the carbonates around the mine areas are characterized by pinkish shading due to iron pigmentation and staining. The paleokarst surface separating the two carbonate sequences of the Naqb Formation is preserved in the ironstone deposits (Figs. 9A, B). The carbonate rocks

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associated with the paleokarst surface usually display Liesegang rings as a result of iron replacement of the carbonate materials (Fig. 9D). Replacive ironstones show a

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variety of fabrics such as colloidal, concretionary, oncolitic-like, brecciated,

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concentric coating box-work and/or reniform aggregates that broadly preserve the

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precursor karstic carbonate structure (Figs. 9E, F). Some dissolution forms, including karst dolines, goethitic pisoids and concentric laminae are cemented and filled with

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Fe, Mn, Si and Ba-bearing minerals (Fig. 9G).

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4.3. Ironstone mineralogy and geochemistry

The principal ore mineralogy of the cherty ironstone deposits consists of

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goethite and hematite minerals, which reach up to 80% in some horizons. Manganese

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minerals, i.e. pyrolusite, jacobsite, todorokite, romanechite, cryptomelane and psilomelane, average 7% in the ironstone. Up to 6% of jacobsite was determined in

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the lowermost part of the ironstone succession associated with iron oxyhydroxides and pyrolusite. Quartz content reaches up to 60% at Ghorabi and El Harra areas. In

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contrast, the quartz content usually does not exceed 5% in ironstone samples from El Gedida except the southern part of the area, where quartz is relatively abundant replacing and/or cementing carbonates (Fig. 1). Barite content is variable, clearly depending on the location where it was observed and easily traced through major faults and stock-work zones. It occurs as big lenses, irregular bodies and open-space cavity or fracture-filling crystals (Fig. 9H). Gangue minerals, mainly apatitefluorapatite, alunite, jarosite, dolomite, minor clay minerals and calcite occur in variable proportions, but do not exceed 7%. The mineralogical analyses of the cherty ironstones were supplemented by whole-rock geochemical analyses (XRF; Table 2) and elemental distribution

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geochemical determinations of the ore-bearing minerals (EMPA; Table 4). These analyses show enrichment of ironstones in trace elements (e.g., Cr, Zn, Cu, As, Rb, S)

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along with FeO and MnO and depletion in Al and Zr (Table 2). The probe analyses of

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these rocks revealed that FeO and MnO contents range between 1.6–82.97% and

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0.06–50.36%, the highest values corresponding to the iron-rich or manganese-rich minerals, respectively. Some Fe/Mn minerals show variable Mg content (up to 13.55%; Table 4), especially in the lowermost part of the ironstone succession where

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presence of jacobsite was determined. CaO content reaches up to 1.64% and

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correlates with MgO, pointing to incorporation of these elements in the ferromanganese minerals. These minerals show variable amount of MnO, FeO, MgO

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and CaO depending on their location. Thus, MgO content in samples from El Gedida

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area is higher than those from Ghorabi and El Harra (Table 4). Presence of BaO was related to Ba-rich manganese oxides (Table 4) as well as barite. SiO2 content is

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mainly related to quartz occurrence rather than to associated clay minerals (Table 4).

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4.4. Ironstone petrography

The petrography of the iron-rich rocks reveals microfacies that are similar to

those shown by the Eocene carbonates but replaced by iron and quartz. In the next paragraphs, the petrography and high-resolution morphometric analyses of the orebearing minerals are described in order of abundance.

4.4.1. Iron oxyhydroxides The iron-rich rocks show iron oxyhydroxides displaying varied morphologies, i.e. amorphous, flakey, acicular, rod-like, tubular, tabular, rosette, fan-like, fibroradiating, botryoidal, and globular micro-fabrics and most commonly preserving

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dolomite pseudomorph fabrics (Fig. 10). These fabrics were recorded as fine groundmass, pore-fillings, corroding quartz crystals and/or cementing dolomite

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pseudomorphs. Frequency and distribution of the main iron oxyhydroxides

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morphologies reflect their relation with the original fabrics of host carbonates. For

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instance, the tubules were recognized in circum-granular cements, the micro-globular fabrics preserved mostly the globular fabrics of the micritized grains, and the tabular and flakey fabrics were recognized in the lowermost part of the succession. Fibro-

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radiating and rosette-like fabrics were observed replacing the highly karstified facies.

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The rhombic dolomite pseudomorphs were recognized in all the transformed Naqb Formation facies.

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The petrographic relationship between the textures and fabrics of carbonates

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and the iron oxyhydroxides show different patterns. The lamination, skeletal fabrics and the concentric layering of the precursor carbonate deposits are preserved after

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replacement and/or cementation by iron oxyhydroxides (Fig. 11A–D). Micritized grains exhibiting micrite envelopes (thickness averaging 0.2 mm; Fig. 11C) and

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micro-pores (Fig. 11D) are usually replaced and/or partly cemented by iron that mostly preserved their precursor microglobular fabrics. The micritized grains are better preserved than their non-micritized counterparts, which suggests that micritization protected the grains from intensive dissolution. As a consequence, the iron oxyhydroxides were observed preferentially replacing the fine-grained carbonates. A variety of karstic features, e.g., concentric carbonates with some colloform quartz, were occasionally preserved (Fig. 11E) and mostly replaced by iron oxyhydroxides (Fig. 11F). Another pattern is represented by extensive pore-filling goethite, which is observed to envelope and/or coat previous iron oxyhydroxides (Fig. 11F).

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4.4.2. Quartz Chert occurrence is more prominent in the ironstone deposits showing features

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similar to the fossiliferous and porous facies F2 and F5 (Figs. 11A, C, G, H). Silica

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occurs as micro-quartz, mega-quartz and/or chalcedony. Quartz is preferentially found

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as pseudomorphs of dolomite crystals and cements infilling the dissolution pores in the carbonate bioclasts, especially the coarse-dolomitized and non-micritized parts (Figs. 10I, 11A, C, G). In the laminated facies, crystalline quartz aggregates occur in-

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between the iron laminae. In general, the small-sized crystalline carbonate precursors

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were preferentially replaced by the iron oxyhydroxides, while the coarse carbonate crystals were transformed into quartz. In addition, silicified grains occur slightly

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corroded by iron.

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Quartz occurs also as fracture-filling and vein-like quartz crystals cutting

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across previous crystalline quartz aggregates and iron oxyhydroxides (Fig. 11H).

4.4.3. Manganese minerals

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The manganese minerals occur in minor amount either mixed with the iron oxyhydroxides groundmass, especially at the lowermost part of the succession (e.g., pyrolusite and jacobsite; Figs. 12A, B), or as late mineral precipitates in pores through all of the succession (e.g., psilomelane, todorokite, romanechite) showing varied morphologies (Figs. 12C, D). Dendritic pyrolusite occurs also filling pores associated with the youngest iron oxyhydroxide phases (Fig. 11F).

4.4.4. Barite Barite occurs as poikilotopic cement, euhedral disseminated crystals, elongated parallel-bedded crystals and/or rosette-like crystals. The poikilotopic barite

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cement consists of optically continuous barite patches up to 5 mm in diameter

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enclosing dolomite molds and relics that are partly replaced by iron (Fig. 11I).

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5.1. Diagenetic sequence and ore mineral paragenesis

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5. Interpretation and discussion

A summary of the sequential arrangement of the diagenetic events in carbonate host rocks and the paragenetic sequence of the ore-bearing minerals

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forming the ironstone in the northern Bahariya is shown in Fig. 13. Time relationships

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and genetic constraints leading to this mineral paragenesis are discussed.

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5.1.1. Micritization

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Micritization processes leading mainly to the formation of micrite envelopes and micro-borings were mainly developed around and inside the skeletal grains

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during the very early diagenetic stages (Figs. 5C, D and 6D). The micritization probably was produced by micro-organisms, e.g., microbes, endolithic algae, fungi,

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bacteria, that bio-eroded the outer part of the carbonate grains by boring small holes, later filled with micrite cement (Adams and MacKenzie, 1998).

5.1.2. Compaction Mechanical compaction resulted in slight re-packing of the skeletal grains making nummulite tests to be oriented under low pressure conditions during early diagenetic stages, as testified by the presence of point and tangential contacts (Fig. 5D). Chemical compaction leading to dissolution of carbonate and formation of fissures and/or stylolites was also observed. The latter compaction features were probably related to the early deformational processes that affected the region

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(Moustafa et al., 2003). The stylolites are clearly observed in the Naqb Formation unlike in the overlying carbonate successions, which suggests that they may have

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been generated during the pre-Middle Eocene deformation phase.

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5.1.3. Dissolution and karst formation

The two karstic phases recognized in the Naqb Formation show variable dissolution features and timing. The first karstification phase took place shortly after

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the deposition of the lower sequence and was then followed by dolomitization

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processes (Figs. 6E, F). The second karstification phase resulted in karst features including the dissolution of dolomites (Figs. 5, 6) and the formation of moldic

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porosities that were later filled by large calcite crystals (Fig. 6F) and/or ore minerals.

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The development of the second karstification phase that affected all the Eocene carbonates probably took place during the Late Eocene-Early Oligocene.

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Under subaerial exposure, epikarst features such as vertical solution pits and shafts were formed in the vadose zone, whilst caves were generated along the phreatic

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water table with the formation of carbonate speleothems and pseudospherulites. The occurrence of nodular and pseudospherulitic calcite fabrics reflects groundwater processes that occurred during the evolution of the paleokarst (Rossi and Cañaveras, 1999; Sanz-Montero, 2009; Hartig et al., 2011). Accordingly, the well-developed subaerial exposure features representing the two karst phases in Eocene carbonates indicate fresh-water diagenesis, including dissolution of carbonate components and development of moldic and vuggy porosities.

5.1.4. Dolomitization

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The dolostones show fair preservation of the depositional fabrics and are characterized by non-uniform grain-size and euhedral to subhedral rhombs that

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typically represent early replacement of predominantly aragonitic limestones (Tucker

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and Wright, 1990). Occurrence of phreatic meteoric cements of dolomite (circum-

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granular cements; Fig. 5B) suggests an early precipitation of dolomite following a sea-level fall. Dolomitization took place after the first karstification phase as indicated by replacement of pseudospherulitic calcite fabrics by dolomite (Figs. 6E, F). This

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pathways for the dolomitizing fluids.

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suggests that the earliest karst processes caused dissolution of the grains and created

Concerning the stable isotope chemistry (Table 2; Fig. 8), δ13CVPDB values for dolomite suggest that carbon derived from marine components was an important

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source for the dolostones (Tucker and Wright, 1990). The negative δ18OVPDB data is indicative of precipitation from fresh water dominated fluids, which is consistent with

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the features of meteoric conditions occurring throughout the carbonate sequence. In addition, the oxygen and carbon isotopes show a covariant trend, which is also

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consistent with dolomitization processes resulting from seawater-fresh water mixing (Budd, 1997). The heavier isotopic compositions of the upper sequence reflect more clear influence of near-normal seawater in the mixing zone (Kyser et al., 2002). The low strontium content of the dolomites (2–73 ppm) determined by probe analyses can be also indicative of mixing zone dolomites (Land, 1973; Land et al., 1975). Both petrographic and geochemical characteristics of the dolomites as well as the sequential stratigraphic context in which they were accumulated, i.e. shallow depositional environments under successive events of sea-level fall and rise, point to the presence of an active meteoric-marine water mixing zone along an oscillating coastline. In this setting, mixing of marine and meteoric water was the most-likely

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triggering mechanism favouring pervasive dolomitization of the primary carbonates

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(Fig. 7).

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5.1.5. Chertification

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Both preservation of dolomite pseudomorphs as quartz and corrosion of quartz crystals by iron oxyhydroxides indicate that silicification was later than dolomite and prior to or contemporaneous with the precipitation of iron oxyhydroxides. Fracture-

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represent a later chertification stage.

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filling and vein-like quartz crystals cutting the quartz crystalline aggregates clearly

Contemporaneous magmatic activity associated with Cenozoic volcanism in

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the region is considered as the most reliable origin for silica-rich fluids replacing and

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cementing the host carbonate rocks. Alternative silica sources such as weathering of the basement rocks or leaching of skeletal siliceous particles in the carbonate are ruled

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out mainly because of the close spatial relation of the chert occurrences to faulted areas and volcanic vents (Fig. 1). Moreover, the primary carbonate host rocks do not

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contain siliceous skeletal grains that could supply silica through leaching processes. Absence of quartz in El Gedida ore deposits is explained by the considerable

distance from the magmatism in the area due to the decrease of geothermal gradient (Holland, 1984). In this setting, Kimberley (1989) stated that the fluids which formed non-cherty iron formations were probably exhaled along observed faults at a lower temperature and pressure than those which formed cherty deposits. In the Algoma type BIF, the silica can be formed through co-precipitation with solid-phase iron minerals (Ewers, 1983), volcanic-derived hydrothermal fluids with contributions from mafic and ultramafic rocks in a marine environment (Frei et al., 2008; Bekker et al., 2010) and/or leaching of basaltic rocks by strongly acidic meteoric water during

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quiescence periods of volcanisms (Zhu et al., 2014). Accordingly, the distribution of silica in the study area, along with the typology of the associated minerals, support the

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relationship between hydrothermal fluids associated with magmatism and the

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mineralization.

5.1.6. Formation of iron oxyhydroxides

The varied fabrics exhibited by goethite and hematite indicate that several

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mechanisms were involved in their formation (Schwertwann and Murad, 1983). The

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amorphous goethite and hematite fabrics reflect rapid precipitation from colloids (Puteanus et al., 1991) while the tabular, tubular, acicular, flakey, and globular fabrics

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indicate slow precipitation rates (Chukhrov et al., 1973; Schwertwann and Murad,

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1983; Afify et al., 2015a). Occurrence of iron oxyhydroxides as rhombic dolomite pseudomorphs proves that they formed later than the dolomites. The preservation of

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dolomite relics, ghosts and pseudomorphs provides evidence that dissolution of carbonate rocks was concomitant with iron oxyhydroxide precipitation (Afify et al.,

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2015a). The occurrence of Liesegang rings and bands as well as box-work structures in ironstones could indicate reducing conditions leading to formation of siderite and/or pyrite that were later oxidized to iron oxyhydroxides (Loope et al., 2011; Kettler et al., 2015). Taking in mind that iron transported by subsurface water in the ferrous state and iron oxides do not precipitate under reducing conditions, formation of other iron-rich minerals would be invoked. Presence of relics and pseudomorphs of dolomite combined without clear recognition of siderite and or pyrite relics indicate that goethite and hematite were most probably precipitated where the ore solution mixed with oxidizing meteoric water. This can be supported also by common occurrence of iron oxyhydroxides filling intraparticle porosity. Variations in dark

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shade and color of the iron oxyhydroxides probably reflect differences in crystal size and degree of dehydration as well as difference in timing of formation (Schwertwann

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and Murad, 1983; Lowe and Byerly, 2007).

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Goethite and hematite can form either by organic or inorganic precipitation.

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The ironstone deposits of the Bahariya area were interpreted as microbial-mediated iron of primary marine origin by Salama et al. (2013). The fair preservation of biotic features such as, stromatolite-like fabrics, oolite cortex, as well as micritized skeletal

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and non-skeletal carbonate grains may be misleading. Occurrence of well-preserved

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micritized fabrics provides additional evidence for the diagenetic origin of iron as micritization usually affects the primary carbonate components. Accordingly, an

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abiotic origin of iron is concluded even though the iron minerals preserve some biotic

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features of their precursor carbonates. The contemporaneity of the iron oxyhydroxides with the silica precipitation as

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well as the co-precipitation of hydrothermal mediated manganese minerals (see discussion, below) supports their hydrothermal origin. Likewise, the geochemistry of

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the iron-rich rocks, when compared with their equivalent carbonate deposits, reflects enrichment in trace elements (e.g., Cr, Zn, Cu, As, Rb, S) and depletion of crustallysourced elements (e.g., Al, Ti, Zr). This supports an authigenic origin for the magmatic hydrothermal fluids and opposes a weathering origin of circulating flows (Nicholson, 1992; Hein et al., 2008; Bekker et al., 2010). Additional iron supply by dissolution of carbonate minerals from the Naqb Formation must be discarded in view of the low Fe content of the dolostones (Table 3).

5.1.7. Formation of manganese minerals

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The association of manganese minerals, e.g., jacobsite and pyrolusite, with the iron oxyhydroxides indicates that these minerals were formed from solutions rich in

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Mn, Mg with Fe that characteristically fractionated to produce high or low Mn/Fe

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ratios, depending on which of the two elements was dominant (Nicholson, 1992; Hein

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et al., 2008; Mohapatra et al., 2009). The presence of jacobsite points to hydrothermal solution where the ex-solution could result in the separation of hematite and jacobsite with decreasing temperature in an oxidizing environment (Nicholson, 1992).

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Magnesium content in ferromanganese minerals may reflect the distance from

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hydrothermal sources. Low Mg contents reflect precipitation from proximal hydrothermal sources, as in Ghorabi and El Harra areas, whilst high calcium values

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reflect a carbonate source (Hein et al., 2008). The formation of manganese minerals as

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acicular, tabular and flakey fabrics, mostly as pore-fillings, and their absence in filamentous fabrics contradict the microbial mediation suggested by Baioumy et al.

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(2013). Due to its higher solubility, manganese was precipitated as pore-filling manganese minerals, especially where the fluids became richer in barium, thus

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reflecting a younger phase of formation of the manganese minerals.

5.1.8. Formation of barite Occurrence of barite in fractures, faults and discontinuities discordant with the bedding clearly indicates that formation of barite was structurally-controlled and most probably related to hydrothermalism in the area (Afify et al., 2015b). Moreover, inclusion of dolomite pseudomorphs partially to totally in the form of hematite within the barite crystals indicate that barite was formed after dolomite and iron oxyhydroxides, respectively. This is also consistent with the formation of barite after quartz. Mixing of sea water and a hydrothermal fluid was suggested as a possible

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origin for the barite formation in the study area by Baioumy (2015). In this setting, barite is commonly precipitated during cooling, especially by mixing of late-stage

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5.2. Cherty ironstone formation model – discussion

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hydrothermal fluids with meteoric water (Rye, 2005).

The petrographic, sedimentological and structural relationships between the iron-rich deposits and carbonate host rocks support a hydrothermal formation model

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for the northern Bahariya cherty ironstones. Diagenetic processes taking place early

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under depositional conditions, i.e. micritization, and later changes throughout compaction, dissolution and dolomitization, slightly modified the main primary

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sedimentary features of the carbonates. These diagenetic changes developed mainly

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under shallow burial conditions. The mineralogical nature of the host rocks played an important role in the transport and geochemical activation of ore fluids.

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Dolomitization and porous sedimentary facies were crucial items in controlling both permeability and porosity, where the dolostone appears to be more porous and

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susceptible to fracturing and brecciation than limestones (Budd and Vacher, 2004). The presence of unaltered claystone beds and laminae could also support selective replacement of carbonates by ore minerals. Our data shows that dolomitization of the Naqb Formation was prior to deposition of the Qazzun Formation during the Late Ypresian whilst the whole set of ore minerals postdated at least the El Hamra Formation (Middle-Upper Eocene) (Fig. 13). This result is consistent with the Late Eocene–Early Oligocene magnetization assigned for the Bahariya area by Odah (2004). In addition, the fluids responsible for the formation of the ore deposits moved throughout major faults that postdate the deposition of the Eocene Naqb, Qazzun and El Hamra formations.

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The proposed model of formation of the ironstone ore deposits from hydrothermal fluids is in agreement with previous interpretation by Dabous (2002),

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Baioumy et al. (2014), Afify et al. (2014, 2015b) and Baioumy (2015). Along this

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line, the lateral association of the ironstone deposits with magmatic rocks south of the

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study area argue for the relationship between the formation of the ironstone and volcanicity. Fault zones played a crucial role in focusing fluid migration into the basin, as can be inferred from the study of many hydrothermal, sediment-hosted ore

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deposits worldwide (Ceriani et al., 2011). The study area shows rejuvenated faults

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that provided a quite suitable extensional setting for hydrothermal driven replacement of the carbonate host rocks. Fracture planes related to the two main fault systems

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behaved as conduits and/or depositional sites for circulating fluids that controlled the

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morphology and distribution of the Fe/Mn, Si and Ba mineralizations. The tectomagmatic approach emphasizes the supply of hydrothermal fluids emanating

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from fractures, deep-seated faults and discontinuities and ended with the replacement of carbonates. Likewise, the distribution of silica, occurrence of jacobsite and

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variations in geochemical compositions with respect to their fluid source location could support this model. The association of jarosite, alunite and halloysite with the ironstone deposits

reflects acid and reduced ore fluids (Rye et al., 1992; Rye, 2005). Under anoxic conditions, iron and manganese can be mobilized in their reduced form (Fe2+ and Mn2+). This fits well with mixing of reducing ore solutions with oxidizing meteoric water circulating through faults and discontinuities (Afify et al., 2015b). Because of the uncommon occurrence of Cenozoic cherty ironstone, the Bahariya ore deposits provide a remarkable perspective to investigate the formation constraints of this kind of rocks. Moreover, they can give some new insight on the

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origin of cherty ironstone of older age, e.g., the Precambrian Algoma type BIF whose

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origin is still a matter of debate (Bekker et al., 2010).

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6. Conclusions

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The cherty ironstone deposits of the Bahariya show lithostratigraphic and sedimentary features similar to those recognized in the Eocene carbonates in which the ironstone is hosted. This strongly supports replacement and cementation of the

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carbonates by silica, mainly quartz, iron oxyhydroxides, manganese-rich and other

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subordinate minerals as a result of late diagenetic and later structurally-controlled processes. The distribution of quartz cements and silica replacements, the presence of

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minerals indicative of formation under hydrothermal conditions, e.g., some

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manganese oxides, and the geochemical enrichment in some elements (Cr, Zn, Cu, As, Rb, S) point to a tectomagmatic genetic model for the cherty ironstones.

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Accordingly, an evolutionary scenario of the ore-forming fluids circulating throughout the regional fault system and sedimentary discontinuities is depicted,

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which is proven by the vicinity of the mineralization to the fractured areas. These processes took place mostly during Late Eocene to Early Oligocene.

Acknowledgments We would like to thank Prof. Dr. Hamdallah A. Wanas (Menoufia University, Egypt), Dr. Emad S. Sallam and Dr. Mohammed K. Zobaa (Benha University, Egypt) for their kind support and help during the field investigation. The authors are indebted to Prof. Brian Jones, Editor-In-Chief, for reviewing and editing of the manuscript and the two reviewers, Profs. David Loope and Cecilio Quesada, for their encouraging comments and annotations that greatly improved an earlier version of the manuscript.

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A standing ovation, in no particular order, goes to Xabier Arroyo Rey, Isabel Gómez Pinilla, Iván Serrano Muñoz, Pedro Lozano and Marián Barajas (Faculty of Geology,

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UCM, Spain) for their help in laboratory studies and mineralogical and chemical

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analyses. Special thanks to Alfredo Fernández Larios (ICTS-CNME Luis Brú of

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Complutense University of Madrid) for microprobe analyses facilities. This work was financially supported by the Egyptian Government in a full fellowship during the study of the first author at Complutense University of Madrid, Spain. This work is a

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part of the activities of Research Groups BSHC UCM-910404 and BSHC UCM-

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910607.

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TABLE CAPTIONS Table 1. Summary of the main sedimentary facies and diagenetic processes

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recognized in the carbonate deposits of the Naqb Formation; column at the right of the

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table focuses on changes of the carbonate rock into cherty ironstone.

Table 2. XRF analyses (Wt. %, ppm) of major oxides and minor elements of 18 bulk samples of carbonates and iron-rich rocks from the Naqb Formation at Ghorabi (Gh)

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and El Harra (Hr). C and O isotope compositions of dolomite samples are listed at the

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lowermost part of the table.

Table 3. Electron microprobe analyses of individual dolomite crystals from different

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carbonate facies of the Naqb Formation (oxides in wt. %). Some analyses represent

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associated minerals, e.g., calcite (point 11), hematite (points 22, 23), quartz (points

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25, 27).

Table 4. Electron microprobe analyses of iron-bearing minerals as well as their

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associated minerals from different ironstone types (oxides in wt. %).

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FIGURE CAPTIONS Fig. 1. A- Location of the study area. B- Geologic map of the northern Bahariya

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Depression (modified after Moustafa et al., 2003) with the main rock units exposed in

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the study area (square in A). C. Location of sections studied at the Ghorabi area. Fig. 2. Geologic profile showing the stratigraphic succession exposed in the northern part of the Bahariya Depression; lithological symbols are the same used in Fig. 1. See

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also legend of Fig. 3 to for symbols of paleokarst and unconformity surfaces.

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Fig. 3. Stratigraphic cross section showing the facies distribution of the Naqb Formation carbonates and their equivalent ironstone deposits at the Ghorabi area. (See

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location of the sections in Fig. 1C).

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Fig. 4. A. Outcrop view of the lower and upper sequences of the Naqb Formation

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separated by a paleokarstic surface (black arrow). The lower sequence is partly replaced by iron (white arrow), while the upper sequence is not replaced. B. Outcrop view of dolostone beds intercalated with marly dolostone (black arrow) at the lower

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part of the Naqb Formation. Note dolostone replaced by ironstone (white arrow). C. Close-up view of stromatolite-like laminated dolostone with white chert laminae and patches. D. Detailed view of the stromatolite-like dolostone where the dolomite laminae are pigmented by reddish iron oxides. E. Small cavities (arrowed) in the lower sequence of the Naqb Formation. F. Outcrop view of sink-holes (white arrow) and a small doline filled by marly deposits (black arrow). G. Dissolution tubes and irregular surface (arrowed) filled by kaolinite clays. H. Close-up view of strongly brecciated dolostone. I. Close-up view of karstic dolostone showing concentric-like arrangement and carbonate speleothems (Scale bar = 10 cm).

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Fig. 5. A. Nummulitic dolostone microfacies. Note selective silicification of skeletal grains and burrows. B. Photomicrograph showing textural features of facies F2

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composed of nummulites, alveolinids, textularids, dascycladacean algae and ooids.

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Note that crystalline dolomite infilling moldic porosity and forms circumgranular

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cements (arrowed) and micrite envelopes around the skeletal grains. C. Fabric characteristic of facies F2 showing some micro-borings. D. Photomicrograph of facies F2 where point-contact and squashed oolitic shapes (arrowed) due to mechanical

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compaction are observed. Mosaics of quartz mesocrystals occur in the dissolution

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pores inside the skeletal and non-skeletal grains as well as in the interparticle porosity. E. Ghost of a benthic foraminifera (textularid; arrowed) in fine-grained euhedral to

D

subhedral dolomite; non-fossiliferous dolostone microfacies F3. F. Loosely-packed,

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medium-grained dolomite rhombs in aggregates of calcite cement (red-stained). All photomicrographs in crossed nicols.

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Fig. 6. A. Crinkled laminate structure formed of fine-grained dolomite with some alternating micro-quartz crystals as observed in the lower part of the picture. B.

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Stromatolite-like laminated dolostone in which laminae of fine-sized dolomite alternate with laminae of quartz (white arrow). The quartz also fills the desiccation cracks (black arrows). C. Stromatolite-like dolostone laminated showing quartz pseudomorphs after intrasedimentary evaporites (white arrows). The quartz also fills the desiccation and bioturbation tubes (black arrows). D. Silicified bivalve dolostone after bivalve packstone microfacies. Note micritization of bioclasts (white arrows). E. Fan-like dolomite. F. Pseudospherulitic dolomite. Note the relics of pseudospherulitic dolomite crystals (first karst phase) inside poikilotopic cements of calcite (second karst phase; arrowed). All photomicrographs in crossed nicols.

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Fig. 7. Interpretative model for the depositional environments of the main facies forming the lower and upper sequences of the Naqb Formation. Inset at the upper part

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of the drawing sketches, the mechanism of dolomitization of the two sedimentary

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sequences in the mixing marine-meteoric zone.

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Fig. 8. Cross plot of stable carbon and oxygen isotope data for the dolomite recorded in the different facies of the Naqb Formation. Note the remarkable linear relation

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between the δ13C and δ18O values.

Fig. 9. A. Outcrop view showing the two sequences formed of carbonate deposits of

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the Naqb Formation. Note the paleokarst surface (arrowed) that separates the two sequences. B. Outcrop view of the ironstone deposit (Ghorabi area) in which a similar

D

lithostratigraphic framework as observed in A can be recognized. C. Stromatolitic-like

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laminated iron-rich rocks overlain directly by bivalve ironstone (black arrow). D.

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Dissolution cavity infilled with ironstone showing Liesegang-ring structure after carbonatic materials embedded by non-replaced kaolinite clays. E. Close-up view of speleogenetic colloidal or reniform aggregates of goethitic ironstone. F. Ironstone

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showing boxwork structure. G. Outcrop view showing a dissolution surface (arrowed) at El Gedida area. H. Fracture-filling barite crystals (white crystals) within the iron ore body. Fig. 10. A–F. SEM photos showing different morphologies of iron oxyhydroxides. A. Rosette-like. B. Fibro-radiating fan-like. C. Tubular, D. Tabular. E. Flakey. F. Acicular and amorphous morphologies (corroding quartz). G. Microglobular fabrics of iron oxyhydroxides typically found in micritized skeletal grain. H. Iron oxyhydroxides in microglobular fabrics preserving the dolomite rhombs of their

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precursor carbonates (white arrow). I. Quartz (Qz) pseudomorphs after dolomite rhombs occurring as relics (white arrows) with some replaced by iron (black arrow).

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Fig. 11. Photomicrographs of A. Oolitic and fossiliferous ironstone where all the

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grains are partly replaced by iron and quartz after dolomite (C.N.). B. Laminated iron

SC R

with high porosity in-between the laminae (PPL). C. Micrite envelopes around bivalve and alveolinid grains replaced by iron (C.N.). D. Micro-borings and micritic cemented

by

iron

on

alveolinid

grain

(PPL).

E.

Carbonates

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envelope

(calcite/dolomite, grey color) with a cockade texture and some colloidal quartz (white

MA

color) (PPL). F. Highly crenulated goethitic iron with pore-filling dendritic manganese oxides (PPL). Note the two generation of iron oxyhydroxides, where the

D

dark ones are of the first generation (white arrows) coated by the second generation of

TE

iron oxyhydroxides (black arrows). G. Nummulitic ironstone with quartz cements in the fossil molds (C.N.). H. A second generation of quartz occurs as fracture-filling

CE P

cement that cuts the first generation of quartz crystals and the iron oxides (C.N.). I. Iron pseudomorphs and dolomite relics are preserved in poikilotopic barite cement

AC

(C.N.). C.N. = crossed nicols, PPL = plane polarized light. Fig. 12. A–D. SEM photos showing different morphologies of manganese minerals. A, B. Acicular jacobsite associated with flakey hematite crystals, C, D. Pore-filling Ba-rich manganese minerals. Fig. 13. Paragenetic sequence of the host carbonate rocks and the ore-bearing minerals in northern Bahariya. Dashed lines indicate minor entity processes whereas shaded continuous lines point to major diagenetic events and products. Red color indicates the minerals found only in the faulted areas.

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Table 1 Petrography

Nummulitic dolostonemarly dolostone (F1)

Indurate, thin dolostone beds, intercalated with marlstoneclaystone. Occurrence of vugs and small caves at the dolostonemarlstone contact. Breccias-like and karstic features. Nummulites and alviolinids dominant, present as molds and/or replaced by calcite and silica. Local bivalve and gastropod shells. Indurate, meter-thick dolostone beds. Massive to strongly brecciated deposits. Slight cross-bedding locally observed though partially erased by dolomitization. Fossil-rich – alveolinids, nummulites, echinoid plates and spines, dasycladacean algae and bivalve shells. Ooids occurring in local patches

Dolomitized mudstone to wackstone, preservation of nummulites and alveolinid tests in finely-crystalline dolomitic groundmass. Biolcasts molds cemented by quartz crystalline aggregates. Dolomite occurs as subhedral, equigranular, vuggy rhombs up to 40 µm maximum size. Scattered quartz and green clay grains. Dolomitized wackestone and packstone to grainstone. Nummulite and alveolinid tests, dasycladacean algae, miliolids, echinoid plates and spines, ooids, rare peliods, cortoids and aggregates floating and/or packed within groundmass of fine to medium-grained, subhedral dolomite rhombs. Moderately to well-sorted ooids exhibit sub-spherical to ovoidal shape. Aggregate of loosely-packed, fine- to medium size dolomite crystals showing scattered tests of diffuse benthic foraminifera (miliolids) and bivalve shells. Dolomite rhombs usually contain cloudy centers. Nonferroan, stoichiometric dolomite. Partial cementation of open-spaces by calcite. Dolomite pseudospherulites. Dolomitized mudstone to wackestone

Massive nonfossiliferous dolostone (F3)

Strongly indurate, meter-thick, massive, bioturbated dolostone. Wide occurrence of irregular concentric laminae and concretions. Small sparse calcite pockets. No fossil remains observable at outcrop scale.

Thick-bedded

Indurate,

meter-thick

massive,

Deposition in shallow subtidal-lagoonal environment.

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MA N

TE D

CE P

AC

Thick-bedded fossiliferous (and oolitic) dolostone (F2)

Interpretation

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Lithology and sedimentary features

CR

Facies

Accumulation in mixed oolitic/bioclastic shoals distributed on a relatively shallow carbonate shelf. Shoals formed under moderate to high energy conditions.

Diagenetic processes and features Micritization; dolomitization; dissolution features– bioclasts molds, local pseudo-spherulitic dolomite, poikilotopic calcite; silicification– bioclasts replaced selectively by quartz. Micritization – affecting both skeletal and nonskeletal particles; dolomitization – in circum-granular cements; dissolution; silicification.

Deposition in intertidal areas prior to subaerial exposure and karstification.

Dolomitization; karstification – development of speleogenic features; local cementation by calcite.

Deposition in shallow

Dolomitization,

Features after replacement by iron Friable to indurate, massive and/or breccias-like black/colored rocks. Bedded structure preserved. Marl/clay intercalations not replaced by iron. Moldic porosity filled by iron and/or silica. Predominance of hematite and minor goethite. Manganiferous iron-rich– presence of pyrolusite and jacobsite. Yellowish to brownish colored rocks. Bedding preserved because of strong silicification filling porosities. Ooids display goethitic and/or hematite concentric layers with relics of dolomite. Replacement of micritized skeletal grains and micrite envelopes by iron oxyhydroxides. Dolomite psuedomorphs replaced by iron and/or silica. Breccias-like rocks – preservation of concentric-like structures related to precursor speleogenic carbonate features. Reniform and botryoidal aggregates of goethitic ironstone. Iron oxyhydroxides display highly porous, irregular colloidal crenulated texture. Acicular goethite infills pores. Strongly brecciated indurate

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breccias-like to bioturbated dolostone. Irregular bed surfaces. Nummulites and some bivalves are common. Concentric structures. Root molds filled by poikilotopic calcite.

with nummulite tests and rare bivalve shells. Aggregates formed of medium-size, euhedral dolomite crystals showing clear outer rims. Stoichiometric dolomite.

subtidal to intertidal environments.

Stromatoliticlike laminated dolostone (F5)

Thinly-laminated dolostone. Planar to wavy and crinkled lamination. Local occurrence of fenestral fabric, bioturbation and root molds. Individual laminae are 3-5 mm thick, grading upwards to thinner laminae. Whitish-colored siliceous laminae increase towards the top, filling discontinuities, desiccation cracks and burrows.

Dolomitized laminated mudstone to wackestone with scattered reworked nummulites and bryozoans remains in a fine crystalline dolomite groundmass. Pseudomorphs of evaporite laminae and nodules replaced by quartz.

Deposition in intertidal to supratidal environments.

Thin-bedded bivalve dolostone (F6)

Very hard, thinly bedded to slightly cross-bedded dolostone. Fossil-rich bivalve shells show local silicification. Usual bioturbations.

Dolomitized bivalve wackestone to packstone showing fine-grained subhedral dolomite rhombs. Bioclasts show either silicification or calcitization.

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MA N

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CR

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fossiliferous dolostone (F4)

Deposition in intertidal-shallow subtidal environment with open circulation and moderate energy conditions.

silicification – both cementation of molds and replacement of grains. Silica content increases towards the upper part of the dolostone beds. Micritization; dolomitization; silicification – millimeter-thick quartz bands; bioclasts molds and vugs are cemented by quartz; dolomite relics within the crystalline quartz aggregates. Dolomitization – in circum-granular cements; local silicification – selective replacement of bivalve shells and burrows; calcitization of skeletal particles.

ironstone– concentric accumulation of iron resulting in oncolitic-like fabrics and box-work structures. Relics of karstified carbonates representing speleogenic masses.

Ironstone deposits show laminated structure that mimic the stromatolite-like fabric of the carbonate. Crinkle and colloform fabrics. Diagonal, inclined and vertical laminations in caves. Locally, the iron-rich laminae alternate with silica bands. Common rosette-like, amorphous and globular fabrics. Bivalve shells occurring as molds preserved and replaced by iron (only in the northeastern part of the Ghorabi area).

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Table 2

Form ula

F1

H r5

G h5

G h8

G h9

F2

G h1 3 F3

G h1 4

G h1 5

G h1 9

F4

G h2 0 F5

G h2 2

G h2 3 F6

Carbonate rocks Minor elements (ppm) 0. 0

2. 2

3. 8

7. 1

8. 4

6. 8

20 .0

10 .7

35 .0

47 .5

Ni

21 .6

11 .1

19 .2

12 .0

24 .9

18 .2

17 .2

17 .6

7. 9

7. 5

10 0. 0 8. 4

Cu

10 .7

7. 4

8. 7

7. 0

9. 3

9. 7

7. 0

7. 4

7. 4

6. 5

Zn

84 .9

28 .3

18 .2

39 .0

17 .4

25 .4

12 .0

0. 7

16 .3

1. 4 1. 2

7. 0 1. 2

3. 1 1. 6

5. 9 1. 5

3. 5 1. 2

2. 5 1. 2

2. 2 1. 7

3. 3 4. 8

65

2

8

37

20

10

6

22 64 17 4 40 5 29 .2 5. 0 40 2

14 4 6. 0 0. 0 5. 0 8. 0 17 06

14 17

99 8

64 7

18 13

9

41

29 6 12 .1

21 0 15 .3

5

Zr Nb Pb S

10 .8

12 .4

0. 7

12 .8

4. 6

3. 5 1. 7

2. 5 1. 2

4. 9 2. 2

4. 9 1. 6

3. 1 1. 8

2

1

1

3

3

21

71

18

85

69

78

12 7

31 5

41

26 6

5

5

6

15

5

75 7 14 0

12 8

5

20

26

83

0

0

0

0

18

95

11

71

10 .9

5. 1

2. 5

2. 6

3. 4

11 .3

13 2 3. 1

42 8. 2 56 .0 1. 4 17 0 54 .0 5. 0 44 .0 2. 0

40 0. 1 21 .4 3. 8

4

40 7. 6 29 .6 1. 4 28 9

6. 1

9. 9

1. 6

22

22

22

22

22

22

22

22

35

22

22

62 9

17 35

47 7

13 28

18 59

53 5

0. 00

0. 00

0. 00

7. 2

7. 00

0. 00

0. 85

0. 85

1. 76

1. 16

0. 85

1. 04

1. 25

6. 97

1. 48

9. 68

0. 00 0. 12 0. 06 10 .0 3 0. 42 0. 47 76 .1 8

0. 00 0. 30 0. 00 23 .7 8 0. 06 0. 27 61 .7 5

0. 0 0. 14 0. 16 12 .1 3 0. 45 4. 06 14 .3 1

0. 00 0. 30 1. 02 31 .0 7 0. 29 0. 17 25 .3 2

0. 00 0. 20 0. 00 19 .3 2 0. 12 0. 81 17 .6 6

NU

MA

22

22

60 5

21 59

10 43

20 71

12 99

19 70

0. 0

0. 00

0. 00

0. 00

8. 91 13 .4 5 0. 49 0. 15 0. 36 18 .0 5 0. 08 0. 27

0. 85 10 .3 0 0. 00 0. 29 0. 19 33 .0 6 0. 19 0. 45

0. 85 15 .0 2 0. 87 0. 27 0. 00 28 .2 8 0. 01 0. 10

20 17 43 74 02 75 4 0 Major elements (%) 61 57 0. 0. .1 .3 00 00 3 9 0. 0. 0. 0. 85 85 85 85 13 15 8. 9. .6 .8 54 69 0 1 0. 1. 0. 2. 13 85 00 05 0. 0. 0. 0. 25 26 08 09 0. 0. 0. 0. 00 00 00 00 26 28 12 14 .4 .3 .3 .6 3 3 3 5 0. 0. 0. 0. 07 08 02 02 0. 0. 0. 0. 17 10 11 07

47 .2 6 0. 85 11 .3 6 2. 47 0. 10 0. 00 17 .6 0 0. 01 0. 11

30 .7 1 0. 85

0. 00 0. 16 0. 00 25 .2 2 0. 04 0. 19

0. 85 10 .0 2 0. 00 0. 31 0. 00 35 .3 2 0. 00 0. 08

2. 21

2. 90

2. 31

1. 77

3. 49

4. 43

1. 58

1. 36 0. 24 1. 00 27 .9 9 0. 22 2. 58

0. 0 0. 09 0. 0

0. 19 0. 66

0. 0

3. 56

1. 87

3

12 .8

22

2. 78

Fe2O

20 .6

22

13 .8

MnO

11 .4

28 6 16 .2

MgO

TiO2

9. 6

36 7 19 .1

0. 85 17 .3 9 0. 56 0. 27 0. 0 30 .0 9 0. 06 0. 16

CaO

7. 3

31 6 9. 3

0. 85

K2O

10 .3

17

Na2O

P2O5

37 .3

14

69 .7 5 0. 85

Al2O3

11 .2

39 2. 4 23 .4 22 .8 6 23 1. 7 29 .8 1. 2

5

0. 94

10 .7

Iron-rich rocks 27 7. 6 10 .5

30

SiO2

G h4 0

22 1

D

Y

G h3 6 F2

23 .1

61 0

14 35

TE

Sr

CE P

Rb

H r1 0

24 .1

AC

Br

G h3 0

G h2 F1

SC R

Cr

As

G h2 4

T

G h1

IP

Samp les

3. 70

4. 23

3. 61

9. 02

19 .6 53 .3

91 .3 42 .2 12 .5 75 .6 19 .3 3. 8

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46 .5 9

43 .7 8

43 .3 8

45 .9 8

δ13C VPD B‰ δ18O VPD B‰

4. 93 7. 71

2. 55 5. 78

1. 72 4. 18

3. 30 5. 99

2. 83 6. 47

1. 05 2. 74

44 45 25 22 25 .7 .0 .6 .4 .5 6 8 6 5 6 C, O isotopes (VPDB ‰) 0. 0. 1. 2. 0. 38 65 52 38 47 3. 4. 1. 0. 1. 31 26 53 77 98

30 .0 5 0. 18 1. 63

45 .0 9

NU MA D TE CE P AC

18 .1 7

24 .8 0

-

-

-

1. 69 4. 33

-

IP

45 .0 2

SC R

L.O.I .

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51

46 .1 5 5. 46 6. 24

32 .0 9

-

-

42 .9 4 2. 34 4. 70

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Al2O3

FeO

MnO

MgO

CaO

Na2O

K2O

TiO2

2.706 2.961 0.626 47.523 3.006 51.794

0.009 0.004 0.070 0.034 0.028 0.039 0.073 0.040 0.055 0.006 0.006 0.037 0.154 0.004 0.074 0.391 0.017 0.058 0.029 0.355 0.369 4.775 0.143 0.079 0.013

0.052 0.257 0.008 0.189 0.389 0.398 0.605 0.690 0.261 0.036 0.063 0.011 0.066 0.014 0.173 0.091 0.176 0.303 0.198 0.459 82.107 82.866 25.367 4.795 1.078 0.275

0.062 0.192 0.003 0.080 0.252 0.066 0.144 0.218 0.070 0.007 0.065 0.028 0.018 0.063 0.074 0.063 0.011 0.032 0.060 0.116 0.028 0.519 0.442 0.037 0.121 0.171 0.087

22.030 21.657 21.616 21.256 21.898 21.4998 21.646 21.131 21.767 21.644 0.553 21.934 22.115 21.696 21.672 21.952 21.626 21.763 21.445 21.734 21.079 1.516 0.312 12.809 10.546 19.858 10.630

30.597 30.661 30.621 30.664 30.439 30.419 30.370 30.145 30.037 30.582 58.917 30.222 30.045 30.466 30.164 30.010 30.689 30.169 30.435 30.355 30.457 1.698 0.825 18.595 14.447 28.508 15.568

0.022 0.026 0.012 0.039 0.031 0.018 0.017 0.023 0.016 0.031 0.022 0.005 0.028 0.012 0.026 0.049 0.038 0.017 0.017 0.008 0.079 0.062 0.045 0.020 0.042

0.001 0.001 0.018 0.004 0.004 0.010 0.014 0.011 0.025 0.011 0.008 0.023 0.003 0.001 0.012 0.025 0.001 0.006 0.011 0.020 0.016

0.008 0.010 0.052 0.054 0.003 -

CR

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MA N

TE D

CE P

P2O5

SO3

BaO

SrO

Total

0.004 0.027 0.009 0.041 0.034 0.008 0.050 0.043 0.074 0.095 0.015 0.003 0.048 0.056 0.070 0.041 0.029 0.070 0.062 0.018 0.065 0.108 0.102 0.073 0.067 0.028

0.018 0.022 0.011 0.008 0.033 0.023 0.002 0.004 0.001 0.003 0.020 0.016 0.003 0.044 0.010 0.014 0.021 0.057 0.005 0.018 -

0.073 0.056 0.053 0.116 0.041 0.044 0.025 0.088 0.055 0.005 0.063 0.019 -

0.159 0.115 0.085 0.139 0.114 0.168 0.095 0.054 0.107 0.035 0.138 0.068 0.053 0.114 0.090 0.072 0.145 0.145 0.173 0.101 0.141 0.054 -

52.945 52.967 52.376 52.340 53.066 52.657 52.727 52.336 52.823 52.816 59.848 52.337 52.266 52.578 52.262 52.437 52.768 52.846 52.644 52.657 52.239 89.171 88.020 62.422 77.718 52.827 78.453

IP

SiO2

AC

Elements Facies Points 1 2 F1 3 4 5 6 7 F2 8 9 10 11 12 13 F3 14 15 16 17 18 19 F4 20 21 22 23 24 F5 25 26 27

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Table 3

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Table 4

6

0.52

0.19

7

11

1.02 2 0.48 6 0.74 1 0.57 7 0.09

0.81 4 0.56 5 1.27 3 3.74 5 0.23

12

0.01

1.01

1.60

13

0.61 8 0.10 6 0.65

0.55 2 0.04 1 0.39

35.7 27 80.4 68 70.2 9 52.4 7 69.1 1 67.3 8 57.5 3 38.8 9 70.0 51 69.8 04 69.2 03 66.7 0 65.0 2 62.1 06 70.0 5 69.8 0 72.2 46 74.2 0

8 9 10

F1 (Ghora bi)

14 15 16

F2 (Ghora bi)

17 18

F4 (Ghora bi)

F5 (Ghora

2.09

1.16 2 1.19 5 1.06 7 -

1.66

19

1.16

20

24

0.00 4 1.08 9 1.35 6 1.50 2 4.77

25

5.43

1.38 0 1.90 1 1.39 8 2.52 3 2.97

26 27

3.51 3 1.09

2.14 9 1.38

28

1.36

1.90

29

1.44 3 1.71

1.38 2 0.36

21 22 23

F3 (El Harra)

0.82

AC

F2 (El Gedida )

0.57

30

K2 O

TiO

P2O

2

5

4.58

11.4 3 11.6 5 13.5 5 11.4 4 11.3 1 0.02

1.30

0.05

0.02

0.03

0.03

1.26

0.00 4 0.01 1 0.10 3 0.52

-

0.01

2.26 2 47.4 12 32.4 68 26.8 61 46.9 1 50.3 6 37.4 39 1.59 1 8.32

0.21 9 0.26 3 0.21 5 0.31 9 6.58

0.13 9 0.31 2 0.24 2 0.29 1 0.44

8.17

5.14 5.29 5.23 8 5.23 0.06

1.41 1.43 4 1.46 0.06

0.00 1 0.01 9 -

SO3

Ba O

Tot al

0.01

0.25

-

0.40

0.02

0.03

0.05 9 -

0.02 6 0.02 8 0.06 1 0.23

0.03 5 0.02

0.12 1 0.55

0.42

0.02 8 0.44 9 0.30 3 0.23 7 0.04

0.17 6 0.23 0 0.15 3 0.14 5 -

0.68 0.26 5 0.44 1 0.50 7 0.22

0.00 4 0.10 2 8.60 7 5.55 4 3.88 9 0.30

54.9 5 56.1 4 57.2 5 54.9 7 56.3 7 84.5 6 75.1 3 83.2 8 82.7 4 79.2 9 60.8 0 62.1 7 83.4 6 83.4 9 80.9 0 78.3 4 74.2 2 72.3 2 61.5 4 55.0 1 75.6 4 75.8 2 75.2 1 79.1 4 78.7 2 70.1 5 75.6 9 76.6 0 75.6 7 77.2 9

T

0.04

Na2 O

0.07

0.03

IP

-

Ca O

-

SC R

5

37.2 8 37.2 6 36.2 4 36.5 5 37.1 2 82.9 7 69.7 86 23.6 86 40.5 88 41.8 88 5.25

Mg O

0.07 5 0.35 9 0.20 4 0.38 3 0.41

0.01 1 0.00 9 0.62 7 0.55 8 0.44 5 0.33

0.21

0.19

0.28

0.03

-

0.12

0.19

0.29 8 0.37 3 0.49

0.30 4 0.08 9 0.29

0.21 2 0.01 7 0.09

0.37 4 -

0.28 3 0.07 0 0.01

0.28 6 0.67 2 0.17

7.01 3 0.03 9 0.20

23.6

0.25

0.24

0.09

0.00 4 0.01

0.35 4 0.02 3 0.03

0.01

0.14

0.11

0.03

0.30

0.04

0.02

0.02

0.23

1.04

0.15

0.03

0.17

0.33

0.03

0.11

0.04

0.23

0.96

0.21

-

0.07

0.28

0.04

0.06

0.02

0.21

10.4 7 0.33 0 0.26 6 0.44 2 3.27

0.70 4 0.10 8 0.10 4 0.12 1 0.31

0.08 4 0.16 5 0.06 0 0.16 4 0.10

-

0.03 7 -

0.21 1 0.03 2 -

-

4.94 2 0.12 2 0.06 3 0.09 6 0.10

0.32

0.89 4 0.09 3 2.31 7 2.15 9 2.13 1 0.89

0.01 7 0.02 6 0.12

0.25 1 0.07 4 0.06 0 0.10 4 0.02

0.14

3.61

0.24

0.01

0.27

0.82

0.03

0.13

0.11 3 0.12 2 0.06

0.54 0 0.33

0.18 9 0.11

1.14 7 2.32

0.11 5 -

0.06 4 0.07

0.10

2.16

0.02

0.60

0.03 5 0.11

0.25 7 0.30

0.08 6 0.04

0.05 7 0.03

0.05 4 0.01 6 0.02 7 -

-

0.27

0.04 5 0.15 7 0.15 6 0.06

0.08 8 0.13

0.03 8 0.00 3

0.02 8 0.00 9

NU

4

0.01 4 0.09 8 -

0.00 1 0.00 7 0.05 8 -

-

Mn O

MA

FeO

D

Al2 O3

2

TE

F1 (El Harra)

SiO

CE P

Elements Facies Poin ts F1 1 (El Gedida 2 ) 3

0.01 6 0.02 7 0.02 0 0.02

0.01

0.02 0 -

0.04 1 0.01 1 0.02 8

ACCEPTED MANUSCRIPT 55

31 32

16.4

1.39

0.03

0.29

1.41

4.70

3.24

0.46

0.21

0.56

0.06

-

1.14 7 0.68 7

0.04 9 -

0.00 8 0.02

-

0.00 5 0.05

0.01

-

-

-

0.04 3

-

-

0.03 7

0.07 6 0.06 5

0.05 1

-

CE P

TE

D

MA

NU

SC R

IP

T

33

49.9 7 97.6 2 96.4 2

AC

bi)

78.7 2 98.9 2 97.3 7

ACCEPTED MANUSCRIPT 56

Highlights 

Well-exposed cherty ironstone is associated with Eocene carbonate in

T

Bahariya area The Eocene carbonates show extensive sea-level related diagenetic features



Ironstone distribution was mainly controlled by post-Eocene major

SC R

IP



faults/fracturing

Fe and Si-rich hydrothermal fluids circulated through the main fault systems



The cherty ironstones show features that can be interpreted as an analog for

NU



AC

CE P

TE

D

MA

BIFs