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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: S0037-0738(15)00203-1
DOI: doi: 10.1016/j.sedgeo.2015.09.010
Reference: SEDGEO 4911
To appear in: Sedimentary Geology
Received date: 3 August 2015
Revised date: 17 September 2015
Accepted date: 18 September 2015
Please cite this article as: Afify, A.M., Sanz-Montero, M.E., Calvo, J.P., Ironstone de-
posits 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 
Afify, A.M.
a,b,*
, Sanz-Montero, M.E.
a
, Calvo, J.P.
a 
a) Department of Petrology and Geochemistry, Faculty of Geological Sciences, Complutense 
University, Madrid, C/ José Antonio Nováis, 2, 28040 Madrid, Spain 
b) Department of Geology, Faculty of Science, Benha University, 13518 Benha, Egypt 
 
* Corresponding author (adelmady@ucm.es) 
ABSTRACT 
This paper gives new insight into the genesis of cherty ironstone deposits. The 
research was centered on well-exposed, unique cherty ironstone mineralization 
associated with Eocene carbonates from the northern part of the Bahariya Depression 
(Egypt). The economically important ironstones occur in the Naqb Formation (Early 
Eocene), which is mainly formed of shallow marine carbonate deposits. Periods of 
lowstand sea-level caused extensive early dissolution (karstification) of the 
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. 
A number of petrographic, sedimentological and structural features including 
the presence of hydrothermal-mediated minerals (e.g., jacobsite), the geochemistry of 
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 
model can contribute to better understanding of the genetic mechanisms of formation 
of banded iron formations (BIFs) that were abundant during the Precambrian. 
 
Keywords: 
Cherty ironstone, dolomitization, tectonic constraints, hydrothermalism, Eocene 
carbonates, Egypt. 
 
1. Introduction 
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 Eocene-
Lower 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 
global warming during the early Paleogene, closely associated with eustatic sea-level 
changes. Recent work by Baioumy et al. (2013, 2014) supports sources and 
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 
exhalite). 
New insight based on field, petrographic, mineralogical and geochemical 
studies of the ironstone deposits in the northern part of the Bahariya Depression (Fig. 
1) is provided in this paper. The similarities between the ironstones and carbonate 
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 
ironstone deposits of Bahariya. Moreover, the scarcity of Phanerozoic, in particular 
Cenozoic cherty ironstone makes the Bahariya ore deposits an interesting case study 
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 
(Late Cenomanian) and Hefuf Formation (Campanian) forming the floor of the 
depression and surrounded by a carbonate plateau of Eocene rocks (El Akkad and 
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 
Formation is overlain by the Qazzun and El Hamra formations, which are poorly 
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 
(Fig. 1). Both the Eocene units and the ironstone deposits are unconformably overlain 
by the Oligocene ferruginous quartzarenite beds of the Radwan Formation (Figs. 1, 
2). Outcropsof basaltic to doleritic igneous dykes, sills, laccoliths and lava flows of 
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 
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 – pre-
Middle 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 
deformation phase continued throughout the Paleocene and Eocene. A similar 
evolution has been documented in the Esna Formation (Paleocene-Early Eocene; 
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 
during the deposition of the Eocene sediments in the Bahariya and Farafra areas (Said 
and Issawi, 1964; Obaidalla et al., 2006). 
The second tectonic phase occurred after the Middle Eocene and before the 
Oligocene. Tectonic inversion continued in the area, leading to the development of 
folds and small domes in the Eocene formations (Said and Issawi, 1964; Moustafa et 
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 
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-
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. 
About 140 samples of carbonate rocks, ironstones, sandstones, clays and other rocks 
were collected. Up to 120 indurated samples were prepared as thin sections and 
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 
(1972) with alizarin red-potassium ferricyanide was used to differentiate between 
carbonate minerals. For high-resolution textural and morphometric analyses, carbon-
coated thin sections and fresh broken pieces were studied using electron microprobe 
analysis (EMPA) and scanning electron microscopy (SEM). A subset of 21 carbon-
coated thin sections were prepared for BSE (backscattered images), SE (secondary 
images) and elemental analyses (in wt. %) on a JEOL Superprobe JXA 8900-M 
wavelength dispersive electron microprobe analyzer (WDS-EMPA) equipped with 
four crystal spectrometers and beam diameter between 2 to 5 µm to minimize damage 
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 
Silicon Drift Detector. Loss on ignition (L.O.I.) was obtained by heating 1g of 
powdered sample at 1000 °C for 1 h. Carbon and oxygen isotopic compositions for 
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 
removed as CO2 after 4 hours of reaction at 25° C. For dolomite, CO2 was extracted 
after an additional 24 hours step at 70° C. 
 
4. Results 
4.1. Carbonate host rocks 
Outcrop observation of the Eocene formations in the northern part of the 
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 
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 
content, the Naqb Formation is Late Ypresian in age (Boukhary et al., 2011). 
The Naqb Formation is subdivided into two sequences separated by a 
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 
underlying and overlying the surface to the Naqb and Qazzun formations, 
respectively. Observations at a regional scale show, however, that the two carbonate 
sequences separated by the paleokarst belong to the Naqb Formation. Brecciation as 
well as formation of speleothems with small- to medium-scaledolines, solution caves 
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 
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 
oolitic) dolostone (F2; Fig. 5B–D) and massive non-fossiliferous dolostone (F3; Figs. 
5E, F). 
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). 
Field and petrographic observations allowed recognition of superimposed 
karstic features pointing to development of two karstification phases in the Naqb 
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 
dolomitized (Figs. 6E, F). Karst features related to the second karstification phase are 
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 
and sub-horizontal dissolution tubes and fractures are either filled by quartz and/or 
large calcite crystals (Fig. 6F). 
The paleoenvironmental interpretation of the carbonate facies of the Naqb 
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 
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 
dolomitization (Figs. 5, 6) which were followed by chertification, formation of iron 
oxyhydroxides, Mn-bearing minerals, and associated gangue minerals in areas where 
the carbonate host rocks were faulted. Timing, paragenesis and relationship of these 
processes are discussed below. 
 
4.1.2. Mineralogy and geochemistry of carbonate host rocks 
The main minerals determined in the carbonate deposits of the Naqb 
Formation are dolomite and quartz, with minor amount of calcite. The dolomite 
content ranges from 30% to more than 90% whereas the calcite content reaches 
locally up to 15%, especially in the non-silicified samples. Quartz content is higher 
where the carbonate deposits approach the iron mineralized beds. 
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 
(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 pore-
filling 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 
2; Fig. 8). Light δ13C (-5.46‰ to -1.05‰) and δ18O values (-7.71‰ to -2.74‰) were 
determined in carbonate samples from the lower sequence of the Naqb Formation 
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). 
 
4.2. Relationships between Eocene carbonate host rocks and ironstones 
Extensive replacement of the Eocene carbonate formations by Fe and Mn 
oxides and quartz occurs in localized areas near major fault lineaments where the 
carbonate rocks are altered partially to totally into cherty ironstone (Table 1; Fig. 3). 
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 
vertical sequential arrangement, thus allowing correlation between the ironstone and 
carbonate deposits and providing evidence for the replacement origin of the former 
(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 
variety of fabrics such as colloidal, concretionary, oncolitic-like, brecciated, 
concentric coating box-work and/or reniform aggregates that broadly preserve the 
precursor karstic carbonate structure (Figs. 9E, F). Some dissolution forms, including 
karst dolines, goethitic pisoids and concentric laminae are cemented and filled with 
Fe, Mn, Si and Ba-bearing minerals (Fig. 9G). 
 
4.3. Ironstonemineralogy and geochemistry 
The principal ore mineralogy of the cherty ironstone deposits consists of 
goethite and hematite minerals, which reach up to 80% in some horizons. Manganese 
minerals, i.e. pyrolusite, jacobsite, todorokite, romanechite, cryptomelane and 
psilomelane, average 7% in the ironstone. Up to 6% of jacobsite was determined in 
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 
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 apatite-
fluorapatite, 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) 
along with FeO and MnO and depletion in Al and Zr (Table 2). The probe analyses of 
these rocks revealed that FeO and MnO contents range between 1.6–82.97% and 
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 
presence of jacobsite was determined. CaO content reaches up to 1.64% and 
correlates with MgO, pointing to incorporation of these elements in the 
ferromanganese minerals. These minerals show variable amount of MnO, FeO, MgO 
and CaO depending on their location. Thus, MgO content in samples from El Gedida 
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 
mainly related to quartz occurrence rather than to associated clay minerals (Table 4). 
 
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 ore-
bearing 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, fibro-
radiating, 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 
pseudomorphs. Frequency and distribution of the main iron oxyhydroxides 
morphologies reflect their relation with the original fabrics of host carbonates. For 
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-
radiating and rosette-like fabrics were observed replacing the highly karstified facies. 
The rhombic dolomite pseudomorphs were recognized in all the transformed Naqb 
Formation facies. 
The petrographic relationship between the textures and fabrics of carbonates 
and the iron oxyhydroxides show different patterns. The lamination, skeletal fabrics 
and the concentric layering of the precursor carbonate deposits are preserved after 
replacement and/or cementation by iron oxyhydroxides (Fig. 11A–D). Micritized 
grains exhibiting micrite envelopes (thickness averaging 0.2 mm; Fig. 11C) and 
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 
similar to the fossiliferous and porous facies F2 and F5 (Figs. 11A, C, G, H). Silica 
occurs as micro-quartz, mega-quartz and/or chalcedony. Quartz is preferentially found 
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-
between the iron laminae. In general, the small-sized crystalline carbonate precursors 
were preferentially replaced by the iron oxyhydroxides, while the coarse carbonate 
crystals were transformed into quartz. In addition, silicified grains occur slightly 
corroded by iron. 
Quartz occurs also as fracture-filling and vein-like quartz crystals cutting 
across previous crystalline quartz aggregates and iron oxyhydroxides (Fig. 11H). 
 
4.4.3. Manganese minerals 
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 
enclosing dolomite molds and relics that are partly replaced by iron (Fig. 11I). 
 
5. Interpretation and discussion 
5.1. Diagenetic sequence and ore mineral paragenesis 
A summary of the sequential arrangement of the diagenetic events in 
carbonate host rocks and the paragenetic sequence of the ore-bearing minerals 
forming the ironstone in the northern Bahariya is shown in Fig. 13. Time relationships 
and genetic constraints leading to this mineral paragenesis are discussed. 
 
5.1.1. Micritization 
Micritization processes leading mainly to the formation of micrite envelopes 
and micro-borings were mainly developed around and inside the skeletal grains 
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, 
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 
been generated during the pre-Middle Eocene deformation phase. 
 
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 
the deposition of the lower sequence and was then followed by dolomitization 
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 
porosities that were later filled by large calcite crystals (Fig. 6F) and/or ore minerals. 
The development of the second karstification phase that affected all the Eocene 
carbonates probably took place during the Late Eocene-Early Oligocene. 
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 
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 
typically represent early replacement of predominantly aragonitic limestones (Tucker 
and Wright, 1990). Occurrence of phreatic meteoric cements of dolomite (circum-
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 
suggests that the earliest karst processes caused dissolution of the grains and created 
pathways for the dolomitizing fluids. 
Concerning the stable isotope chemistry (Table 2; Fig. 8), δ13CVPDB values for 
dolomite suggest that carbon derived from marine components was an important 
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 
the features of meteoric conditions occurring throughout the carbonate sequence. In 
addition, the oxygen and carbon isotopes show a covariant trend, which is also 
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 
(Fig. 7). 
 
5.1.5. Chertification 
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-
filling and vein-like quartz crystals cutting the quartz crystalline aggregates clearly 
represent a later chertification stage. 
Contemporaneous magmatic activity associated with Cenozoic volcanism in 
the region is considered as the most reliable origin for silica-rich fluids replacing and 
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 
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 
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 
relationship between hydrothermal fluids associated with magmatism and the 
mineralization. 
 
5.1.6. Formation of iron oxyhydroxides 
The varied fabrics exhibited by goethite and hematite indicate that several 
mechanisms were involved in their formation (Schwertwann and Murad, 1983). The 
amorphous goethite and hematite fabrics reflect rapid precipitation from colloids 
(Puteanus et al., 1991) while the tabular, tubular, acicular, flakey, and globular fabrics 
indicate slow precipitation rates (Chukhrov et al., 1973; Schwertwann and Murad, 
1983; Afify et al., 2015a). Occurrence of iron oxyhydroxides as rhombic dolomite 
pseudomorphs proves that they formed later than the dolomites. The preservation of 
dolomite relics, ghosts and pseudomorphs provides evidence that dissolution of 
carbonate rocks was concomitant with iron oxyhydroxide precipitation (Afify et al.,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 
and Murad, 1983; Lowe and Byerly, 2007). 
Goethite and hematite can form either by organic or inorganic precipitation. 
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 
and non-skeletal carbonate grains may be misleading. Occurrence of well-preserved 
micritized fabrics provides additional evidence for the diagenetic origin of iron as 
micritization usually affects the primary carbonate components. Accordingly, an 
abiotic origin of iron is concluded even though the iron minerals preserve some biotic 
features of their precursor carbonates. 
The contemporaneity of the iron oxyhydroxides with the silica precipitation as 
well as the co-precipitation of hydrothermal mediated manganese minerals (see 
discussion, below) supports their hydrothermal origin. Likewise, the geochemistry of 
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 crustally-
sourced 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 
Mn, Mg with Fe that characteristically fractionated to produce high or low Mn/Fe 
ratios, depending on which of the two elements was dominant (Nicholson, 1992; Hein 
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). 
Magnesium content in ferromanganese minerals may reflect the distance from 
hydrothermal sources. Low Mg contents reflect precipitation from proximal 
hydrothermal sources, as in Ghorabi and El Harra areas, whilst high calcium values 
reflect a carbonate source (Hein et al., 2008). The formation of manganese minerals as 
acicular, tabular and flakey fabrics, mostly as pore-fillings, and their absence in 
filamentous fabrics contradict the microbial mediation suggested by Baioumy et al. 
(2013). Due to its higher solubility, manganese was precipitated as pore-filling 
manganese minerals, especially where the fluids became richer in barium, thus 
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 
hydrothermal fluids with meteoric water (Rye, 2005). 
 
5.2. Cherty ironstone formation model – discussion 
The petrographic, sedimentological and structural relationships between the 
iron-rich deposits and carbonate host rocks support a hydrothermal formation model 
for the northern Bahariya cherty ironstones. Diagenetic processes taking place early 
under depositional conditions, i.e. micritization, and later changes throughout 
compaction, dissolution and dolomitization, slightly modified the main primary 
sedimentary features of the carbonates. These diagenetic changes developed mainly 
under shallow burial conditions. The mineralogical nature of the host rocks played an 
important role in the transport and geochemical activation of ore fluids. 
Dolomitization and porous sedimentary facies were crucial items in controlling both 
permeability and porosity, where the dolostone appears to be more porous and 
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), 
Baioumy et al. (2014), Afify et al. (2014, 2015b) and Baioumy (2015). Along this 
line, the lateral association of the ironstone deposits with magmatic rocks south of the 
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 
deposits worldwide (Ceriani et al., 2011). The study area shows rejuvenated faults 
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 
behaved as conduits and/or depositional sites for circulating fluids that controlled the 
morphology and distribution of the Fe/Mn, Si and Ba mineralizations. The 
tectomagmaticapproach emphasizes the supply of hydrothermal fluids emanating 
from fractures, deep-seated faults and discontinuities and ended with the replacement 
of carbonates. Likewise, the distribution of silica, occurrence of jacobsite and 
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 (Fe
2+
 and 
Mn
2+
). 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 
origin is still a matter of debate (Bekker et al., 2010). 
 
6. Conclusions 
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 
carbonates by silica, mainly quartz, iron oxyhydroxides, manganese-rich and other 
subordinate minerals as a result of late diagenetic and later structurally-controlled 
processes. The distribution of quartz cements and silica replacements, the presence of 
minerals indicative of formation under hydrothermal conditions, e.g., some 
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. 
Accordingly, an evolutionary scenario of the ore-forming fluids circulating 
throughout the regional fault system and sedimentary discontinuities is depicted, 
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, 
UCM, Spain) for their help in laboratory studies and mineralogical and chemical 
analyses. Special thanks to Alfredo Fernández Larios (ICTS-CNME Luis Brú of 
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 
part of the activities of Research Groups BSHC UCM-910404 and BSHC UCM-
910607. 
 
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TABLE CAPTIONS 
Table 1. Summary of the main sedimentary facies and diagenetic processes 
recognized in the carbonate deposits of the Naqb Formation; column at the right of the 
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) 
and El Harra (Hr). C and O isotope compositions of dolomite samples are listed at the 
lowermost part of the table. 
Table 3. Electron microprobe analyses of individual dolomite crystals from different 
carbonate facies of the Naqb Formation (oxides in wt. %). Some analyses represent 
associated minerals, e.g., calcite (point 11), hematite (points 22, 23), quartz (points 
25, 27). 
Table 4. Electron microprobe analyses of iron-bearing minerals as well as their 
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 
Depression (modified after Moustafa et al., 2003) with the main rock units exposed in 
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 
also legend of Fig. 3 to for symbols of paleokarst and unconformity surfaces. 
Fig. 3. Stratigraphic cross section showing the facies distribution of the Naqb 
Formation carbonates and their equivalent ironstone deposits at the Ghorabi area. (See 
location of the sections in Fig. 1C). 
Fig. 4. A. Outcrop view of the lower and upper sequences of the Naqb Formation 
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 
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 
composed of nummulites, alveolinids, textularids, dascycladacean algae and ooids. 
Note that crystalline dolomite infilling moldic porosity and forms circumgranular 
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 
compaction are observed. Mosaics of quartz mesocrystals occur in the dissolution 
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 
subhedral dolomite; non-fossiliferous dolostone microfacies F3. F. Loosely-packed, 
medium-grained dolomite rhombs in aggregates of calcite cement (red-stained). All 
photomicrographs in crossed nicols. 
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. 
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 
of the drawing sketches, the mechanism of dolomitization of the two sedimentary 
sequences in the mixing marine-meteoric zone. 
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 
between the δ13C and δ18O values. 
Fig. 9. A. Outcrop view showing the two sequences formed of carbonate deposits of 
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 
lithostratigraphic framework as observed in A can be recognized. C. Stromatolitic-like 
laminated iron-rich rocks overlain directly by bivalve ironstone (black arrow). D. 
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 
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). 
Fig. 11. Photomicrographs of A. Oolitic and fossiliferous ironstone where all the 
grains are partly replaced by iron and quartz after dolomite (C.N.). B. Laminated iron 
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 
envelope cemented by iron on alveolinid grain (PPL). E. Carbonates 
(calcite/dolomite, grey color) with a cockade texture and some colloidal quartz (white 
color) (PPL). F. Highly crenulated goethitic iron with pore-filling dendritic 
manganese oxides (PPL). Note the two generation of iron oxyhydroxides, where the 
dark ones are of the first generation (white arrows) coated by the second generation of 
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 
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 
(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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Fig. 1 
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Fig. 8 
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Fig. 9 
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Fig. 10 
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Fig. 11 
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Fig. 13 
 
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Table 1 
Facies Lithology and sedimentary features 
 
Petrography Interpretation Diagenetic processes 
and features 
Features after replacement by iron 
Nummulitic 
dolostone-
marly 
dolostone (F1) 
Indurate, thin dolostone beds, 
intercalated with marlstone-
claystone. Occurrence of vugs and 
small caves at the dolostone-
marlstone 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. 
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. 
Deposition in shallow 
subtidal-lagoonal 
environment. 
Micritization; 
dolomitization; 
dissolution features–
bioclasts molds, local 
pseudo-spherulitic 
dolomite, 
poikilotopic calcite; 
silicification– 
bioclasts replaced 
selectively by quartz. 
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. 
Thick-bedded 
fossiliferous 
(and oolitic) 
dolostone 
(F2) 
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 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. 
Accumulation in 
mixed 
oolitic/bioclastic 
shoals distributed on 
a relatively shallow 
carbonate shelf. 
Shoals formed under 
moderate to high 
energy conditions. 
Micritization – 
affecting both 
skeletal and non-
skeletal particles; 
dolomitization – in 
circum-granular 
cements; dissolution; 
silicification. 
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. 
Massive non-
fossiliferous 
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. 
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. Non-
ferroan, stoichiometric dolomite. 
Partial cementation of open-spaces by 
calcite. Dolomite pseudospherulites. 
Deposition in 
intertidal areas prior 
to subaerial exposure 
and karstification. 
Dolomitization; 
karstification – 
development of 
speleogenic features; 
local cementation by 
calcite. 
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. 
Thick-bedded Indurate, meter-thick massive, Dolomitized mudstone to wackestone Deposition in shallow Dolomitization, Strongly brecciated indurate 
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fossiliferous 
dolostone (F4) 
 
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. 
silicification – both 
cementation of molds 
and replacement of 
grains. Silica content 
increases towards the 
upper part of the 
dolostone beds. 
ironstone– concentric accumulation 
of iron resulting in oncolitic-like 
fabrics and box-work structures. 
Relics of karstified carbonates 
representing speleogenic masses. 
Stromatolitic-
like 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. 
Micritization; 
dolomitization; 
silicification – 
millimeter-thick 
quartz bands; 
bioclasts molds and 
vugs are cemented by 
quartz; dolomite 
relics within the 
crystalline quartz 
aggregates. 
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. 
Thin-bedded 
bivalve 
dolostone 
(F6) 
Very hard, thinly bedded to slightly 
cross-bedded dolostone. Fossil-rich 
bivalve