Possibility of Mixed Origin of Rare Earth Elements in Sedimentary Marine Apatites: A Case Study from Phosphorites in the Cretaceous (Campanian-Maastrichtian) Duwi Formation, Abu-Tartur Plateau, Egypt

Introduction

It is well known that the rare earth elements (REEs) incorporated in the crystal lattice of sedimentary apatite are commonly derived from ambient seawater. This study documents, for the first time, the possibility of mixed origin of apatite REEs present in the Egyptian Western Desert phosphorites, known as the Abu-Tartur phosphorites, one of the most important accumulations of sedimentary phosphorites in the Middle East and North Africa. Shale-like patterns of REEs, negative Ce anomalies, and a (La/Sm)_N⁻(La/Yb)_N binary diagram of the studied phosphorites indicate that the incorporation of REEs into apatite crystal lattices has occurred from the ambient seawater by substitution during late-stage diagenesis. The second origin is attributed to REE-bearing supergene fluids, which resulted in the occurrence of sedimentary britholite as green rims and patches containing about 14.93 wt% total REEs in places where the black phosphorites are gradually oxidized into the brown variety. For instance, the intensive chemical weathering induces the crystal lattice of carbonate fluorapatite to preferentially release Ca²⁺ and CO₃²⁻ ion species to solution, resulting in the formation of a carbonate-depleted layer in which REEs, particularly heavy REEs, are incorporated into the preferential Ca²⁺ sites inside the altered apatite lattice, leading to britholite formation.

Rare earth elements (REEs) are distinguishable from the other elements by their little fractionation during the sedimentary cycle. So the recorded changes in REE distribution patterns can be employed as a useful geochemical means to study and reconstruct paleoenvironmental conditions in oceans and estuaries as well as geological evolutionary processes in general (Deng et al. [16]). REEs are mainly brought to seawater by eolian and riverine fluxes (Sholkovitz et al. [54]; Chen et al. [12]) and then removed at the mixing zone under low-salinity conditions by salt-induced, flocculated, iron-rich, organic colloids in the order of light REEs (LREEs), medium REEs (MREEs), and then heavy REEs (HREEs; Haley et al. [27]; Johannesson et al. [34]), resulting in the enrichment of HREEs in seawater relative to their counterparts. Existing research has shown that seawater is considered to be the only source of REE content of marine apatites, resulting in a negative Ce anomaly, a seawater-like REE pattern, and enrichment in HREEs relative to LREEs (Fleischer and Altschuler [24]; McArthur and Walsh [43]; Wright et al. [60]; Shields and Stille [53]; Sa et al. [51]). However, seawater-derived patterns of REEs of marine phosphorites enriched in LREEs over HREEs are also recorded (e.g., Baioumy [6]; Kechiched et al. [35]). This finding is attributed to postdepositional changes during diagenesis and/or the mobilization of HREEs relative to LREEs under intensive chemical weathering (Shields and Stille [53]; Baioumy [6]).

The incorporation of REEs into the crystal lattice of marine apatite results in two types of patterns: HREE-enriched patterns that are largely similar to those of the epicontinental seawater and MREE-enriched patterns that are concave downward (Reynard et al. [48]). Generally, REEs are taken up by apatites through direct and indirect mechanisms. The former occurs through one of the following reaction processes: (1) quantitative uptake from the ambient seawater, (2) preferential adsorption of LREEs on apatite crystal faces controlled by crystal-chemical properties, and (3) late diagenetic substitution of MREEs (Reynard et al. [48]; Lécuyer et al. [38]). The indirect mechanism is promoted by the adsorption of REEs on active particles of Fe- and Mn-oxyhydroxides and then returned to the pore water under reducing conditions; hence, they are diagenetically incorporated into the crystal lattice of apatite (De Baar et al. [15]). Igneous apatites acquire REEs primarily from hydrothermal fluids during a coupled substitution ( ${\mathrm{Ca}}^{\mathrm{+2}}+{\mathrm{P}}^{\mathrm{+5}}={\mathrm{REEs}}^{\mathrm{+3}}+{\mathrm{Si}}^{\mathrm{+2}}$ ), leading to formation of a REE-bearing phosphate variety called britholite whose occurrence in apatite-rich rocks is considered pronounced evidence of hydrothermal activity (Lira and Ripley [39]; Uher et al. [58]). At this point, the present study sheds light on the possibility of britholite formation in sedimentary marine apatite deposits and the possibility of mixed origin of apatite REEs. The Egyptian Western Desert phosphorites, called Abu-Tartur phosphorites, have been selected as a case study because of their high REE content compared with counterparts in the Nile Valley and Red Sea districts (El-Kammar and El-Kammar [22]). Combinations of macroscopic, microscopic, mineralogical, and geochemical investigations of the Egyptian phosphorites are integrated to assess the different sources of apatite REEs and the paleoenvironmental conditions under which the incorporation of REEs occurs.

Geology and Stratigraphy

The studied Egyptian phosphorites cover approximately 1200 km2 associated with other rock units in the form of a suboval plateau, called the Abu-Tartur plateau, which is located between latitudes 25°10′N and 25°26′N and longitudes 29°51′E and 30°10′E along the road between the Kharga and Dakhla oases (fig. 1a). The ground surface of the Abu-Tartur plateau is completely covered by the Nubia plains composed of fluviatile to marginal marine sedimentary rocks and represents rapid and sequential periods of marine regression and transgression (fig. 1b). The oldest rock unit consists of fluviatile tabular and trough cross-bedded sandstone of the Late Jurassic Six Hills Formation, which is more or less conformably overlain by clay and shale sediments of the Aptian Abu Ballas Formation. The latter represent a transitional state from the fluvial deposition to the beginning of marine deposition (Klitzsch [37]; Hermina [28]). However, the marine deposition was incised at the inception of the Albian age, during which white kaolinitic sandstone sediments of the Sabaya Formation accumulated. From the Late Cenomanian to the Early Turonian, the sedimentary basin was occupied by marginal marine claystone alternating with siltstone and sandstone of the Maghrabi Formation (Klitzsch [37]; Barthel and Hermann-Degen [9]). The marine deposition was interrupted again during the Early Turonian, resulting in accumulation of massive, laminated, and cross-bedded sandstone of the Middle Turonian (Barthel and Boettcher [8]). The Campanian age is considered the beginning of marine transgression following the tectonically controlled Early Turonian regression, leading to the accumulation of variegated shale deposits of the Quseir Formation (Hermina [28]). Phosphate sediments were deposited at high rates along with intercalations of black shale and glauconitic sandstone during the Late Cretaceous relative sea level rise. These sediment accumulations are represented by the Duwi Formation, which is assigned a Campanian-Maastrichtian age (Youssef [62]; Issawi [31]). The occurrence of sand-sized phosphorite mostly accumulated in the lower part of the Duwi Formation, whereas the middle is dominated by black shale, above which glauconitic sandstone with thin layers of evaporites occurs. The Campanian-Maastrichtian succession is terminated by a thin layer, about 40 cm thick, of nodular phosphorite. The cooccurrence of sand-sized phosphorite with black shale, followed by glauconitic sandstone, reveals a sequential change from a high-energy and wave action environment to a low-energy marine environment. Moreover, the coursing-upward sequence at the top implies a short time of relative sea level fall, which is rapidly followed by an open marine regression during the Maastrichtian age, resulting in the accumulation of shale, marl, and interbedded limestone of the Dakhla Formation (Ahmad et al. [1]; El Ayyat [17]). The latter is unconformably overlain by a reef-like limestone with shale intercalations deposited in a shallow shelf environment and assigned to the Early Paleocene Kurkur Formation. The shallow flooding of seawater in the Early Paleocene gave way to a deeper sea in the Late Paleocene, resulting in the deposition of the Garra Formation during the Late Paleocene–Early Eocene. The Garra Formation consists of well-bedded, massive white limestone separated from the underlying Kurkur limestone by a clay bed marker. Since the retreat of seawater and the uplifting phase in the post–Early Eocene time, terrestrial conditions have prevailed over the Kharga-Farafra stretch, resulting in the deposition of sand dunes, playa deposits, and gravel fills (Hermina [28]).

MAP: Figure 1. Geological map of the Abu-Tartur plateau in the Western Desert of Egypt (after Hermina [28]; El Ayyat [17]; a), with Landsat images showing the Abu-Tartur plateau (b) and the eastern and western mining sectors where phosphate samples were collected (c).

Material and Methods

Thirteen representative phosphate samples were collected from two measured stratigraphic sections at the Abu-Tartur phosphate mine (fig. 1c). The first one is located at the eastern sector of the mining area at latitude 25°25′33.3N between longitudes 30°05′16.6E and 30°05′14.7E. This section measures 10.5 m thick, of which 8 m is recorded as the Campanian-Maastrichtian Duwi Formation; the remaining thickness of 2.5 m is assigned to the Maastrichtian Dakhla Formation (fig. 2a). The lower part of the Duwi Formation, where most of the samples were obtained, represents sandy, coarse-grained, moderately sorted, and moderately hard phosphate deposits occurring as one bed and containing thin layers of black shale stained with iron oxyhydroxides and intersected by microveinlets of evaporites. The sandy phosphate bed is covered by intercalations 3 m thick of brownish- to dark-gray flaky shale, thin layers of evaporites, and glauconite-rich facies. This part is capped by nodular calcareous phosphorite about 10 cm thick, which serves as a lithological boundary between the Campanian-Maastrichtian Duwi Formation and the Maastrichtian Dakhla Formation (Tantawy et al. [57]). The remaining part of the studied section consists mainly of calcareous shale assigned to the Maastrichtian Dakhla Formation. On the western side of the Abu-Tartur mine, phosphate deposits belonging to the Duwi Formation are recorded at latitude 25°24′30.47N and longitude 30°03′50.03E at the lower part of a section about 40 m thick (fig. 2b). The studied samples were collected from sandy, hard, medium- to coarse-grained, and moderately sorted phosphate deposits about 5 m thick. The studied section is occupied by brown phosphate with a lens of black phosphate and intercalations of black fissile shale. Phosphate deposits are overlain by brownish- gray shale 12 m thick alternating with yellowish-green glauconite deposits. These clayey deposits are overlain by 13 m of blackish- to light-gray shale alternating with brownish marl. The top 10 m of the studied section is reported to be 6.5 m of brown to black shale deposits of the Dakhla Formation and 3.5 m of load casting limestone of the Kurkur Formation. It is worth mentioning that the lower 5 m containing phosphate deposits is the only accessible part from which we obtained samples. The measured sections are drawn with sampling horizons using SedLog 3.1 software (fig. 3).

Graph: Figure 2. Field view of phosphate deposits on the eastern (a) and western (b) sides of the Abu-Tartur mine, showing the shade variations of phosphate deposits as well as the associated rock units.

Graph: Figure 3. Measured stratigraphic sections of phosphate deposits on the eastern (a) and western (b) sectors of the Abu-Tartur mining area. Sand: vf = very fine; f = fine; m = medium; c = coarse; vc = very coarse. Gravel: gran = granular; pebb = pebble; cobb = cobble; boul = boulder.

For microscopic investigations, thin sections were prepared and studied by a polarized light microscope (Olympus-BX51). The bulk mineralogy was determined using a X-ray diffractometer (XRD) with monochromatic Cu kα radiation ( $\lambda =1.540$ Å, 40 KV, 25 mA). The obtained patterns were interpreted using the X'Pert HighScore Plus software. The mineral composition was emphasized by Fourier transform infrared spectroscopy (FTIR). For major, trace, and rare earth elements, the studied samples were analyzed using X-ray fluorescence (XRF) and inductively coupled plasma (ICP) mass spectrometry hosted at Acme Lab, Canada. The ICP sample preparation followed the analysis procedures mentioned by El-Habaak et al. ([18]).

Results

Macroscopic Investigations

The studied hand specimens of Abu-Tartur phosphorites are characterized by noticeable color variations, including black, gray, and brown, due to the gradual oxidation of organic matter and associated pyrite (El-Kammar and Basta [21]). The gradual weathering is manifested by the irregularly shaped gray patches inside the black phosphorite (fig. 4a, 4b). Under prolonged chemical weathering, the latter was reduced to black spots surrounded by the gray phosphorite (fig. 4c); moreover, iron oxyhydroxides started to appear (fig. 4d). One of the most attractive features is the occasional appearance of irregularly shaped green halos and rims in places where black phosphorite is transformed into a gray variety (fig. 4e, 4f). The green variety is also gradually oxidized into brown phosphorite (fig. 4g, 4h) and completely disappears from the highly oxidized samples.

Graph: Figure 4. Macroscopic investigations of Abu-Tartur phosphorites, showing the gradual oxidation of black phosphorite, passing through the gray variety, into the brown phosphorite (a–d). Green rims and spots are observed through the oxidation zones (e, f). These rims are also oxidized (g; arrows) and occasionally occur as residual spots inside of the brown phosphorite (h; arrows).

Microscopic Investigations

Under the microscope, the phosphate-bearing components include oval to suboval, angular to subangular, and rarely rounded structureless phosphate pellets, coprolites, teeth, and scales cemented by cryptocrystalline carbonate fluorapatite (fig. 5a). Phosphate pellets vary between fine (127 μm) and medium (570 μm) sand size and can be classified as moderately sorted pellets, whereas poorly sorted counterparts (180 μm–1.5 mm in diameter) are well observed in the uppermost part of the studied sections. The main difference between the nonoxidized and oxidized phosphate types is the color of the pellets, which is deep black in the case of nonoxidized pellets because of the high content of organic matter (fig. 5b). Under intensive chemical weathering, the organic matter is oxidized and gradually leaves brownish pellets behind. The transition from nonoxidized to oxidized ore types occurs through successive steps. This phenomenon was studied by preparing two thin sections from the contact between the black phosphorite and the gray variety. The deep black pellets are gradually transformed into gray pellets. The latter are invaded by brownish coloration inward from the grain periphery (fig. 5c, 5d). Furthermore, the prolonged intensive weathering has resulted in white-colored pellets containing brown spots and rims. This coloration can be attributed to the removal of carbonate from the crystal structure of carbonate fluorapatite, leading to the formation of white-colored fluorapatite (fig. 5a, 5e). This kind of transformation has been discussed by McClellan and Van Kauwenbergh ([44]) and Van Kauwenbergh et al. ([59]). In places, aggregates of pellets, teeth, and scales resembling rock fragments are also observed (fig. 5f), which may indicate the reworked origin from authigenic phosphorites as discussed by Baioumy and Tada ([7]). Carbonate minerals are represented here by calcite and ankerite. The former serves as an epigenetic intergranular cement containing traces of cryptocrystalline carbonate fluorapatite (fig. 6a), whereas the latter occurs in the uppermost part of the studied section as well-developed rhombic crystals associated with angular to subangular, poorly sorted phosphate pellets (fig. 6b). In all cases, carbonate minerals and phosphate components are corroded and replaced by gypsum (fig. 6c, 6d). Quartz is less abundant than carbonate and sulfate minerals and appears as angular to subangular detrital grains varying in diameter between 50 and 200 μm. Pyrite is observed in the nonoxidized ore type only as open space–filling materials (fig. 6e). Hematite appears at the expense of pyrite in the oxidized ore as deep red irregular patches diffused through some interstitial spaces and sometimes as fine cluster aggregates inside phosphate pellets (fig. 6f). The last words in this section are related to the occurrence of green material, which is thought to be britholite. It is randomly distributed throughout phosphate deposits and found more in the nonoxidized phosphate than in the oxidized counterpart. Morphologically, britholite sometimes occurs as an intergranular diffused material corroding both phosphate pellets and cryptocrystalline carbonate fluorapatite-rich cement (fig. 7a–7c). The complete britholitization of phosphate pellets is also observed in the nonoxidized ore type (fig. 7d). Irregularly shaped grains, along with patches and stains of britholite, are perceived through phosphate pellets, teeth, and scale fragments (fig. 7e–7h). The common occurrence of deep-green pellets in the nonoxidized phosphorite, along with the appearance of yellowish-green pellets in the oxidized ore type, indicates that the britholitization was formed early in the nonoxidized phosphorite and then gradually eliminated under intensive chemical weathering.

Graph: Figure 5. Photomicrographs of Abu-Tartur phosphorites, showing oval to suboval and angular to subangular phosphate pellets along with fish teeth and scales cemented by collophane (reddish brown in color) in the oxidized (a) and nonoxidized (b) phosphorites as well as rock fragment–like aggregates of phosphate pellets and fish scales set in collophane-rich cement (f). The gradual elimination of organic matter (c, d) and the transformation of francolite into white spots of fluorapatite (a, e) are also observed.

Graph: Figure 6. Photomicrographs of Abu-Tartur phosphorites, showing the occurrence of nonphosphate components as calcite-rich cement (a), well-developed rhombic crystals of ankerite (b), gypsum-rich cement containing traces of calcite and collophane (c, d), black pyrite as open space–filling materials, and clusters of hematite/goethite inside phosphate pellets (f).

Graph: Figure 7. Photomicrographs of Abu-Tartur phosphorites, showing the occurrence of green britholite as an intergranular cement material (a–c) and suboval pellets (d). Under intensive chemical weathering, the britholitization is gradually eliminated (c–g), resulting in brownish-green pellets in the oxidized ore type (h).

Mineralogy

The bulk mineralogy was investigated for black, gray, and brown phosphorites as well as the green patches on the eastern and western sides of Abu-Tartur mine. The phosphate-bearing minerals are represented in all samples by fluorapatite, carbonate fluorapatite, and britholite (figs. 8–11). Pyrite and montmorillonite are detected only in the black and gray phosphate types, whereas hematite and goethite appear in the brown ore type as an alteration product of these minerals. The high diffraction intensities of gypsum, compared with calcite and pyrite, along with microscopic investigations, support the claim of El-Kammar and Basta ([21]) that gypsum is considered to be an alteration product of calcite and pyrite. A trace amount of quartz is well observed in all samples. Ankerite appears only in the brown phosphorite located directly beneath the black shale–rich facies with diffraction intensities higher than that of fluorapatite (fig. 11b). In addition to XRD, FTIR analysis was performed for the green patches and rim to confirm the occurrence of britholite. The obtained infrared (IR) pattern (fig. 12) was interpreted according to the standard patterns of Chukanov ([13], p. 382–423). The vibrational absorption bands of britholite occupy most of the IR spectral range and appear at 3410, 1684, 1620, 1456, and 1041 cm−1. Fluorapatite and gypsum are also detected through the IR spectral range. Fluorapatite occurs at wave numbers 1041, 670, 603, and 471 cm−1, whereas gypsum absorbs only at 3545 and 116 cm–1.

Graph: Figure 8. X-ray diffractometer patterns of black phosphorites on the eastern (a) and western sides (b) of the Abu-Tartur mine.

Graph: Figure 9. X-ray diffractometer patterns of gray phosphorites on the eastern (a) and western sides (b) of the Abu-Tartur mine.

Graph: Figure 10. X-ray diffractometer patterns of brown phosphorites on the eastern (a) and western sides (b) of the Abu-Tartur mine.

Graph: Figure 11. X-ray diffractometer patterns of the green patches and rims (a) and the brown phosphorite located directly beneath the black shale–rich facies (b).

Graph: Figure 12. Fourier transform infrared spectroscopy pattern of the green patches and rims emphasizes the occurrence of britholite.

Geochemistry

The distribution of the major oxides of the Abu-Tartur phosphorites is listed in table A1. The P2O5 contents vary between 14.67 and 22.45 wt%, with a noticeable enrichment in the oxidized samples compared with the nonoxidized ones. The strong positive correlation ( $r=0.61$ ) between P2O5 and CaO (fig. 13a) is due to the occurrence of calcium ion species in the crystal lattice of apatite (Awadalla [4]). The CaO/P2O5 ratio ranges between 1.73 and 2.58, indicating the carbonate mineral association with phosphorites (Slansky [55]). The apatite crystal lattice is considered to be an open lattice allowing a great number of substitutions, such as the substitution of (SO4)–2 for (PO4)–3 (Nathan [47]), resulting in strong negative correlations between these ion species (fig. 13b). The relationship between P2O5 and SiO2 is commonly negative because of the aforementioned substitution. However, the occurrence of detrital mineral phases such as clay minerals results in a positive correlation between P2O5 and SiO2 (fig. 13c; Dar and Khan [14]). The Al2O3 also shows a weak positive relationship with P2O5 (fig. 13d) related to the ionic substitutions in the apatite crystal lattice or the adsorbed alumina in clay minerals rich in P2O5 (Khan et al. [36]). The positive correlation between F and P2O5 (fig. 13e) indicates the incorporation of F into the apatite crystal lattice (Awadalla [4]). The F/P2O5 weight ratio is used to discriminate between the highly substituted carbonate fluorapatite (0.148) and fluorapatite (0.089), as discussed by McClellan and Van Kauwenbergh ([44]). By comparing the studied samples (table A1), it is clear that there is a transformation from carbonate fluorapatite into fluorapatite. The strong positive correlation between Fe2O3 and SO4 ( $r=0.74$ ) can be ascribed to the admixed pyrite (fig. 13f).

Graph: Figure 13. Variation of P2O5 with CaO, SO4, SiO2, Al2O3, and F, as well as the relation between Fe2O3 and SO4.

Regarding trace elements (table A2), their occurrence in phosphorites is mostly owing to the intensive microbial degradation of the associated organic matter, resulting in the liberation of trace elements into the ambient pore water and then their incorporation into the carbonate fluorapatite crystal lattice (Jarvis [33]). This is revealed by the negative correlations of Ni, Pb, Hg, and Co with P2O5 (fig. 14a–14d) due to the ionic substitutions inside of apatite crystal lattice (Howard and Hough [29]). The absence of a noticeable relationship of Sr with P2O5 and CaO ( $r=0.04$ ), along with the positive correlation between Sr and SO4 (fig. 14e), can explain the mild enrichment of Sr, which is liberated from the diagenetic calcite cement and then accommodated with the weathering product (sulfates; El-Kammar and Basta [21]). The positive correlation between U and P2O5 (fig. 14f) results from the ionic substitution of U for Ca in the apatite crystal lattice due to similar ionic sizes (Awadalla [4]). The studied phosphorites are characterized by U/Th ratios varying between 3.26 and 87.92, which are somewhat close to the U/Th ratios of the Miocene phosphorites of the California basin (1.7–101). Moreover, this ratio can reach up to 300, as in the Miocene phosphorites from the Chatham Rise (Baturin [9]). The plot of the U/Th ratios for the studied samples through and close to the hydrothermal field of the U/Th diagrams of Boström et al. ([11]) and Rona ([49]) does not imply a hydrothermal contribution to such deposits (fig. 15).

Graph: Figure 14. Variation of P2O5 with Ni, Pb, Hg, Co, and U, as well as the relation between SO4 and Sr.

Graph: Figure 15. Log U/log Th binary diagram of Boström et al. ([11]) and Rona ([49]) showing different hydrothermal fields of various sediments: I = Trans-Atlantic Geotraverse (TAG) hydrothermal area; II = Galapagos spreading-center deposits; III = amphitrite hydrothermal sediments; IV = Red Sea hot brine deposits; V = East Pacific Rise crest deposits; VI = Langban hydrothermal sediments; VII = ordinary manganese nodules; VIII = ordinary pelagic sediments; IX = laterites; X = fossil hydrothermal deposit. It is clear that the Abu-Tartur phosphorites fall through and near the hydrothermal field of the East Pacific Rise crest deposits.

The studied phosphorites contain considerable concentrations of REEs, varying between 269.5 and 929.9 ppm (table A3). This range is close to that of the Chinese REE-bearing clay deposit type, which contains from 500 to 2000 ppm total REEs (ΣREEs; Wu et al. [61]). The average REE content of the nonoxidized phosphate samples (767.02 ppm) is higher than that of the oxidized ones (712.4 ppm). However, some oxidized phosphorites seem to be relatively enriched in REEs, with contents varying between 899.9 and 929.9 ppm. This can be attributed to the adsorption of REEs on iron oxyhydroxides particles, which formed by phosphorite weathering (Gnandi and Tobschall [26]). The studied phosphorites are characterized by negative Ce/Ce* (average 0.79) and positive Eu/Eu* (average 1.22) and composed of marine apatites, as shown in the binary diagram of El-Kammar ([20]), which discriminates between marine and terrestrial apatites by plotting the LREEs/HREEs ratio against ΣREEs (fig. 16). There is LREE enrichment and HREE depletion, with an average 8.9 ΣLREEs/ΣHREEs ratio. This is clearly manifested by the concave-down North American shale composite (NASC)-normalized LREE pattern. Likewise, the studied phosphorites are also enriched in the MREEs, as illustrated by the bell-shaped normalized pattern (fig. 17). The fractionation ratio (La/Yb)N between LREEs and HREEs is somewhat constant for all samples, implying little or no fractionation between LREEs and HREEs. However, the (La/Sm)N ratio is estimated at values slightly increased, from 0.78 to 1.30. The plot of phosphorite samples in the (La/Sm)N−(La/Yb)N binary diagram of Reynard et al. ([48]) falls close to the substitution mechanism (fig. 18). The hydrogenic and hydrothermal effects on oceanic sediments and nodules are reflected by the Ce/La ratio. For instance, the least oxygenated conditions of ocean bottom currents are characterized by a Ce/La ratio of >2 (Zabel and Schulz [63]; Shatrov and Voitsekhovskii [52]), whereas the hydrothermal sedimentary deposits are distinguished by a Ce/La ratio of <2 (Liu et al. [40]). For the present study, the Ce/La ratio is estimated at an average of 2.23.

Graph: Figure 16. Plot of light/heavy rare earth elements (LREEs/HREEs) against the sum of REEs to discriminate between terrestrial and marine apatites (after El-Kammar [20]).

Graph: Figure 17. Seawater-like NASC-normalized rare earth elements (REE) pattern of Abu-Tartur phosphorites showing a concave-up pattern of light and medium REEs (arrows).

Graph: Figure 18. The (La/Sm)N−(La/Yb)N binary diagram of Reynard et al. ([48]) illustrates that the studied samples fall close to the substitution mechanism during late diagenesis.

Discussion

Origin of Phosphorites

The structureless pellets, along with the fractured fish teeth and scales as well as the rock fragment–like aggregates of pellets and bioclasts, indicate the reworked origin of the Abu-Tartur phosphorites (Baioumy and Tada [7]; Baioumy 2007). The abundance of fish teeth and scales suggests that the studied phosphorites were reworked from a high-productivity environment (Baioumy [5]). The authigenic source of phosphorite was not far from the depositional basin as indicated by the low sphericity degree and moderate sorting. The occurrence of collophane as an intergranular cement material also implies redeposition in an inner-shelf environment (Al-Bassam et al. [2]).

Origin of REEs

As mentioned, the studied phosphorites contain considerable concentrations of REEs. reaching up to 929.9 ppm. The comparison of the NASC-normalized REE patterns of the Abu-Tartur phosphorites with those of McArthur and Walsh ([43]) shows that the Abu-Tartur phosphorites have a shale-like pattern, which suggests that the REE enrichment was diagenetically derived from the high input of terrigenous material (El-Haddad and Ahmed [19]; Awadalla [4]; Baioumy [6]) or from the associated pyrite, which serves as REE scavenger from the ambient seawater (Ismael [30]). Although pyrite and montmorillonite are completely absent in the oxidized samples (figs. 10, 11b), the REE enrichment is still present. So pyrite and montmorillonite cannot be the main reason for the REE enrichment in the Abu-Tartur phosphorites.

The present study explains the REE incorporation into the Abu-Tartur phosphorites through the following three mechanisms, which have been proposed and explained by Reynard et al. ([48]) and Lécuyer et al. ([38]). The first one is represented by a quantitative uptake of REEs from the ambient seawater ("hydrogenous origin"), resulting in a seawater-like REE pattern. The ability of apatite to fix REEs in its crystal structure, along with the physicochemical similarities between REEs (Baioumy [6]), leads to a quantitative uptake without fractionation among the REEs. The second one takes place during early diagenesis, where LREEs are preferentially adsorbed on apatite crystal surfaces over HREEs. Consequently, the (La/Yb)N fractionation ratio drastically increases, whereas the (La/Sm)N fractionation ratio will be little affected. The third mechanism includes the substitution of REEs for Ca+2 sites in the apatite crystal lattice during late diagenetic recrystallization, in which the (La/Sm)N fractionation ratio should increase over the (La/Yb)N fractionation ratio. By comparison with the present study, the plot of phosphorite samples through the (La/Sm)N−(La/Yb)N binary diagram of Reynard et al. ([48]) shows that the incorporation of REEs into the apatite crystal lattice was done by the substitution mechanism during late diagenesis. Such a substitution type is preferred for the MREEs, along with Ce and Nd, because of the similar ionic radii (Wright et al. [60]; Morad and Felitsyn [46]). This is supported by the bell-shaped normalized pattern at Ce and Gd (fig. 16). However, the shale-like pattern, along with negative Ce anomalies, postulates that the REEs were directly incorporated from the ambient seawater without fractionation under prevailing oxic conditions (McArthur and Walsh [43]). This controversy can be solved, depending on the reworked origin of the Abu-Tartur phosphorites, which supposes that the incorporation of REEs from seawater has not been originally accomplished by substitution. According to Ruttenberg and Berner ([50]), apatites are common minerals in organic-rich marine sediments such as phosphorites and form at or near the sediment-water interface, where phosphates are diffused into pore water via the microbial degradation of organic matter. Such authigenic apatites quantitatively absorb REEs from seawater without fractionation immediately after deposition. In this case, the resultant REE patterns are enriched in HREEs, like seawater (Shields and Stille [53]). On reworking, the incorporation mechanisms of REEs in reworked apatites are derived by adsorption or substitution, depending on the dominant fractionation ratio, such as (La/Yb)N or (La/Sm)N. So the reworked marine apatites, as in the present case study, are supposed to acquire their REE contents during authegensis and redeposition processes.

Until now, the aforementioned data refer to the seawater origin of REEs in the studied marine apatites. However, the green rims and patches observed in the Abu-Tartur phosphorites suggest another origin. Both XRD (fig. 11a) and FTIR (fig. 12) assays show that the dominant composition of these green patches is assigned to britholite. Also, XRF analysis of such patches reveals high concentrations of REEs reaching up to 14.93 wt% ƩREEs (table A4), which can support this suggestion. Britholite is the REE-bearing phosphate mineral that results from a coupled substitution of REE3+ and Si4+ for P5+ and Ca2+ in the apatite crystal lattice and commonly occurs in the igneous intrusions as a result of hydrothermal alteration of REE-bearing minerals (e.g., Lira and Ripley [39]; Arden and Halden [3]; Uher et al. [58]). However, the occurrence of britholite in the Abu-Tartur phosphorites is not associated with veins, veinlets, or even void fillings that characterize hydrothermal REE-bearing mineralization (Mariano [41]). The studied phosphorites lack geological evidence to convincingly support the occurrence of hydrothermal activities associated with phosphate deposition. Even the positive Eu anomaly reported here (1.19–1.26) is not necessarily an indication for high-temperature hydrothermal activities (Sverjensky [56]; Shields and Stille [53]; Garnit et al. [25]), but instead it indicates an anoxic diagenetic environment, as in the case of the Sonrai phosphorites in India (Khan et al. [36]). The REE-bearing phosphate minerals (e.g., apatite and monazite) may have formed through magmatic, hydrothermal, or supergene (weathering-type) processes (Mariano [41], [42]; Mgonde [45]). Because of the lack of evidence for hydrothermal activities here, along with the occurrence of britholite in the oxidation zone of black phosphorites (fig. 4e–4h), britholite is suggested to have formed by REE-bearing supergene fluids (fig. 19) that interacted with the weathered apatite. The calculated Ʃ(La-Eu)/Ʃ(Gd-Lu) ratio ranges between 4.56 and 11.46, indicating that the Abu-Tartur phosphorites have been subjected to intensive chemical weathering under humid conditions (Shatrov and Voitsekhovskii [52]). The formation of britholite in the studied marine apatites can be explained, depending on the transformation of carbonate fluorapatite into fluorapatite, which is inferred from the F/P2O5 weight ratio average of 0.099 (McClellan and Van Kauwenbergh [44]). For instance, the carbonate fluorapatite becomes metastable under intensive chemical weathering, resulting in a preferential release of Ca2+ and CO32– ion species from the apatite crystal lattice to solution relative to phosphate, and hence a carbonate-depleted layer may form around the weathered surface of apatite, with a tendency to fluorapatite formation (Jahnke [32]). The Ca2+ sites in the apatite crystal lattice are mostly preferred for REE substitution (Fleet and Pan [23]). The weathering of apatite enables REEs to substitute for Ca2+ in the apatite crystal lattice (Mgonde [45]). Accordingly, it is proposed that the britholite formation in the studied marine apatites has stemmed from the reincorporation of REEs into Ca2+ sites present in the aforementioned carbonate-depleted layer under extreme chemical weathering. The detected britholite is enriched in HREEs, which are more mobile than LREEs on weathering (Shields and Stille [53]). This scenario sheds more light on the reasons for the postdepositional enrichment of REEs in the Abu-Tartur phosphorites suggested by Baioumy ([6]).

Graph: Figure 19. Simplified drawing illustrates the mixed origin of rare earth elements (REEs) in sedimentary marine apatites from seawater (a) and supergene enrichment (b). MREEs = medium rare earth elements.

Conclusion

As implications of the present study, REEs incorporated into the crystal lattice of marine apatites could be derived from two different sources. The first one is represented by the ambient seawater, resulting in a shale-like pattern associated with a negative Ce anomaly. The second source is supergene enrichment under extreme chemical weathering, during which REEs, in particular HREEs, are incorporated into the weathered crystal lattice of apatite, resulting in green rims and patches of britholite in places where the black phosphorites are gradually oxidized. This is the first documentation of britholite appearance in sedimentary phosphorites without the commonly associated hydrothermal activities.

Acknowledgment

We are grateful for the efforts made by M. Abdel-Galil, chairman of the Abu-Tartur phosphate mines, during the collection of phosphorite samples.

Appendix

Graph

Table A1. Distribution of Major Oxides (wt%) of the Studied Phosphorite Samples

Elements	Black phosphorites, nonoxidized variety	Brownish-gray phosphorites, oxidized variety
P.3	P.4	P.5	P.12	P.1	P.2	P.6	P.7	P.8	P.9	P.10	P.11	P.13
P2O5	14.67	15.01	14.87	15.36	19.92	19.86	19.27	22.45	19.87	20.38	18.03	17.62	17.26
SiO2	5.96	5.04	5.19	4.17	5.02	6.36	6.39	11.11	7.64	5.37	2.06	3.87	5.83
Al2O3	.89	1.31	1.56	.63	1.26	2.08	2.31	.79	2.0	1.28	ND	1.43	1.8
CaO	30.44	33.05	33.64	32.77	38.76	34.49	35.13	38.86	36.59	38.93	42.56	40.40	34.50
MgO	.55	.59	1.02	.69	1.02	1.0	1.34	1.0	1.01	1.13	4.28	6.19	1.27
SO4	20.01	22.19	18.05	19.79	8.13	9.73	11.92	6.21	11.39	12.86	7.93	5.62	14.97
K2O	.01	.09	ND	.01	ND	.41	.14	ND	.22	.06	ND	ND	.08
Fe2O3	8.76	8.03	9.56	8.09	4.35	6.14	5.76	4.65	4.11	3.88	4.67	5.25	4.54
F	1.75	1.49	1.57	1.94	1.81	1.78	1.92	1.8	1.84	2.0	1.76	1.83	1.6
LOI	15	12	14	15	18	16	15	13	15	13	18	17	15
Sum	98.04	98.8	99.46	98.45	98.27	97.85	99.18	99.87	99.67	98.89	99.29	99.21	96.85
F/P2O5	.119	.099	.105	.126	.090	.088	.099	.080	.092	.098	.097	.103	.093

1 LOI = loss on ignition; P.3–P.5 = black phosphorite at the eastern sector of Abu-Tartur mine; P.12 = black phosphorite at the western sector of Abu-Tartur mine; P.1, P.6, P.7 = brownish-gray phosphorite at the eastern sector of Abu-Tartur mine; P.8–P.11, P.13 = brownish-gray phosphorite at the western sector of Abu-Tartur mine; ND = not determined.

Graph

Table A2. Distribution of Trace Elements (ppm) of the Studied Phosphorite Samples

Elements	Black phosphorites, nonoxidized variety	Brownish-gray phosphorites, oxidized variety
P.3	P.4	P.5	P.12	P.1	P.2	P.6	P.7	P.8	P.9	P.10	P.11	P.13
Ni	71	74.6	59.2	76.5	49.8	8	1.9	8.4	12.8	9.4	22.1	19.4	7.4
Pb	20.19	19.69	18.69	19.45	17.44	20.4	19.69	18.69	4.74	30.98	10.71	9.25	19.45
Co	16.3	30.7	16.7	16.6	8.9	2	1.2	2.4	6.2	3.7	8.4	7.2	8.1
Cu	25.28	24.43	20.23	20.80	12.19	18.49	14.80	15.17	12.61	13.21	7.65	7.08	11.89
Hg	183	153	172	265	55	143	153	172	223	95	25	22	265
Sr	1520	1620	1340	1730	1500	1420	1480	1200	1790	2140	2150	1850	1900
U	24.4	25.5	24.7	24.6	28.1	29.7	25.5	24.7	123.1	26.7	20.3	18.8	24.6
Th	5.6	5.9	5.5	6.1	8.6	7.7	5.9	5.5	1.4	7.2	3.5	2.3	6.1
U/Th	4.35	4.32	4.49	4.03	3.26	3.85	4.32	4.49	87.92	3.7	5.8	8.17	4.03

2 P.3–P.5 = black phosphorite at the eastern sector of Abu-Tartur mine; P.12 = black phosphorite at the western sector of Abu-Tartur mine; P.1, P.6, P.7 = brownish-gray phosphorite at the eastern sector of Abu-Tartur mine; P.8–P.11, P.13 = brownish-gray phosphorite at the western sector of Abu-Tartur mine.

Graph

Table A3. Distribution of Rare Earth Elements (ppm) of the Studied Phosphorite Samples

Elements	Black phosphorites, nonoxidized variety	Brownish-gray phosphorites, oxidized variety
P.3	P.4	P.5	P.12	P.1	P.2	P.6	P.7	P.8	P.9	P.10	P.11	P.13
La	136.9	141	141.9	142.7	185.3	171.1	173.2	36.9	165.3	203.3	96.4	71.7	148.5
Ce	345	284	309	252	250	317	278	170	223	243	270	227	244
Pr	32.46	35.08	34.61	35.03	49.11	44.91	43.58	6.76	41.85	49.11	22.49	16.75	33.39
Nd	138	147	144	144	204	184	180	27.69	167.8	210	94.79	70.21	141
Sm	28.05	31.14	29.15	30.45	42.37	38.86	37.45	5.09	35.44	43.89	19.59	14.25	29.43
Eu	8.48	9.17	8.49	8.67	12	11.44	10.98	1.44	11	13.14	5.77	4	8.72
Gd	33.25	33.57	31.49	32.79	45.84	42.85	40.25	6.15	39.47	47.86	23.04	15.49	30.98
Tb	4.63	4.91	4.79	4.73	6.66	6.27	5.95	.87	5.83	7.02	3.41	2.29	4.56
Dy	29.19	30.31	29.10	29.19	41.05	38.69	36.10	5.43	35.47	45.82	20.42	14.08	28.71
Ho	6.5	6.87	6.57	6.68	9.54	8.32	8.14	1.29	8.27	10.08	4.66	3.24	6.55
Er	17.48	18.12	18.01	17.87	25.02	23.12	21.57	3.68	20.93	25.93	12.41	8.67	17.59
Tm	2.56	2.60	2.58	2.56	3.53	3.25	3.03	.53	3.02	3.78	1.81	1.27	2.44
Yb	15.49	16.86	15.89	15.50	21.45	19.70	18.88	3.12	18.89	22.93	10.99	7.97	15.64
Lu	2.48	2.60	2.49	2.55	3.31	2.98	2.96	.56	3.03	3.80	1.73	1.30	2.54
ƩREEs	800.7	763.8	778.7	724.9	899.9	912.6	860.1	269.5	779.3	929.9	587.5	458.2	714.7
ƩLREEs	722.4	681.4	699.2	645.8	789.3	810.2	763.5	254	683.9	810	532.1	419.4	636.7
ƩHREEs	78.33	82.27	79.43	79.08	110.5	102.3	96.63	15.48	95.44	119.3	55.43	38.82	78.03
(La/Yb)N	.77	.73	.78	.8	.75	.76	.80	1.01	.76	.77	.76	.78	.83
(La/Sm)N	.88	.81	.87	.84	.78	.79	.83	1.30	.84	.83	.88	.90	.91
Ce/Ce*	.88	.91	.67	.8	.59	.82	.72	1.44	.6	.55	1.31	1.14	.78
Eu/Eu*	1.22	1.24	1.23	1.2	1.19	1.23	1.24	1.13	1.29	1.25	1.19	1.18	1.26
Ce/La	3	2	2	2	1	2	2	5	1	1	3	3	2
Ʃ(La–Eu) /Ʃ(Gd–Lu)	6.17	5.59	6.01	5.48	4.75	5.28	5.28	11.46	4.77	4.56	6.48	7.43	5.55

3 P.3–P.5 = black phosphorite at the eastern sector of Abu-Tartur mine; P.12 = black phosphorite at the western sector of Abu-Tartur mine; P.1, P.6, P.7 = brownish gray phosphorite at the eastern sector of Abu-Tartur mine; P.8–P.11, P.13 = brownish gray phosphorite at the western sector of Abu-Tartur mine. Ʃ denotes sum; LREEs = light REEs; HREEs = heavy REEs.

Graph

Table A4. Distribution of Rare Earth Elements and Selected Major Oxides of the Green Rims in the Studied Phosphorites

	P2O5	SiO2	CaO	SO4	Pr2O3	Nd2O3	Sm2O3	Eu2O3	Gd2O3	Tb2O3	Dy2O3	Ho2O3	Yb2O3	LOI
wt%	6.73	9.80	34.20	10.87	.051	.361	1.532	2.487	1.525	.715	5.321	1.202	1.740	18

4 LOI = loss on ignition.

1 Author for correspondence; email: mahmoud.sabry@sci.svu.edu.eg. Ahmad, F.; Farouk, S.; and Abd El-Moghny, M. W. 2014. A regional stratigraphic correlation for the upper Campanian phosphorites and associated rocks in Egypt and Jordan. Proc. Geol. Assoc. 125:419–431. 2 Al-Bassam, K.; Aba-Hussain, A. A.; Mohamed, I. Q.; and Al-Rawi, Y.T. 2010. Petrographic classification of east Mediterranean phosphorite deposits. Iraqi Bull. Geol. Mining 6:59–79. 3 Arden, K., and Halden, N. M. 1999. Crystallization and alteration history of britholite in rare-earth-element-enriched pegmatitic segregations associated with the Eden Lake Complex, Manitoba, Canada. Can. Mineral. 37:1239–1253. 4 Awadalla, G. S. 2010. Geochemistry and microprobe investigations of Abu-Tartur REE-bearing phosphorite, Western Desert, Egypt. J. Afr. Earth Sci. 57:431–443. 5 Baioumy, H. M. 2007. Iron-phosphorous relationship in the iron and phosphorite ores of Egypt. Chem. Erde Geochem. 67:229–239. 6 ———. 2011. Rare earth elements and sulfur and strontium isotopes of upper Cretaceous phosphorites in Egypt. Cretaceous Res. 32:368–377. 7 Baioumy, H. M., and Tada, R. 2005. Origin of Late Cretaceous phosphorites in Egypt. Cretaceous Res. 26:261–275. 8 Barthel, K. W., and Boettcher, R. 1978. Abu Ballas Formation (Tithonian/Berriasian; Southwestern Desert, Egypt), a significant lithostratigraphic unit of the former "Nubian Series." Mitt. Bayer. Staatsslg. Palaeont. Hist. Geol. 18:154–166. 9 Barthel, K. W., and Hermann-Degen, W. 1981. Late Cretaceous and Early Tertiary stratigraphy in the Great Sand Sea and its SE margins (Farafra and Dakhla Oasis), SW Desert, Egypt. Mitt. Bayer. Staatsslg. Palaeont. Hist. Geol. 21:141–182. Baturin, G. N. 1982. Phosphorites on the seafloor: origin, composition and distribution. Amsterdam, Elsevier Science, 345 p. Boström, K.; Rydell, H.; and Joensuu, O. 1979. Langban: an exhalative sedimentary deposits? Econ. Geol. 74:1002–1011. Chen, J.; Algeo, T. J.; Zhao, L.; Chen, Z. Q.; Cao, L.; Zhang, L.; and Li, Y. 2015. Diagenetic uptake of rare earth elements by bioapatite, with an example from Lower Triassic conodonts of South China. Earth-Sci. Rev. 149:181–202. Chukanov, N. V. 2014. Infrared spectra of mineral species. Dordrecht, Springer. Dar, S. A., and Khan, K. F. 2016. Depositional environment of phosphorites of the Sonrai Basin, Lalitpur District, Uttar Pradesh, India. In Marghany, M., ed. Applied studies of coastal and marine environments. London, InTech, p. 840–845. De Baar, H. J. W.; German, C. R.; Elderfield, H.; and van Gaans, P. 1988. Rare-earth element distributions in anoxic waters of the Cariaco Trench. Geochim. Cosmochim. Acta 52:1203–1219. Deng, Y.; Ren, J.; Guo, Q.; Cao, J.; Wang, H.; and Liu, C. 2017. Rare earth element geochemistry characteristics of seawater and porewater from deep sea in western Pacific. Sci. Rep. 7:16539. doi:10.1038/s41598-017-16379-1. El Ayyat, A. M. 2015. Lithostratigraphy, sedimentology, and cyclicity of the Duwi Formation (Late Cretaceous) at Abu Tartur plateau, Western Desert of Egypt: evidences for reworking and redeposition. Arab. J. Geosci. 89:99–124. El-Habaak, G.; Askalany, M.; Galal, M.; and Abdel-Hakeem, M. 2016.Upper Eocene glauconites from the Bahariya depression: an evidence for the marine regression in Egypt. J. Afr. Earth Sci. 117:1–11. El-Haddad, M. A., and Ahmed, E. A. 1991. Facies control on the distribution of some trace and rare earth elements in Egyptian phosphorites. J. Afr. Earth Sci. 12:429–435. El-Kammar, A. 1985. Geochemistry and REE patterns for terrestrial and marine skeletal apatites from Egypt. J. Jpn. Assoc. Min. 80:21–33. El-Kammar, A. M., and Basta, E. Z. 1983. Chemical weathering of the economic phosphates of Abu-Tartur, Western Desert, Egypt. Chem. Geol. 38:321–328. El-Kammar, A. M., and El-Kammar, M. 2002. On the trace elements composition of the Egyptian phosphorites: a new approach. In International Conference on Geology of the Arab World, 6th (Cairo University, February), p. 227–244. Fleet, M. E., and Pan, Y. 1997. Site preference of rare earth elements in fluorapatite: binary (LREE+HREE)-substituted crystals. Am. Mineral. 82:870–877. Fleischer, M., and Altschuler, Z. S. 1982. The lanthanides and yttrium in minerals of the apatite group: a review. U.S. Geol. Surv. Open-File Rep. 82–783, 11 p. Garnit, H.; Bouhlel, S.; Barca, D.; and Chtara, C. 2012. Application of LA-ICP-MS to sedimentary phosphatic particles from Tunisian phosphorite deposits: insights from trace elements and REE into paleo-depositional environments. Chem. Erde Geochem. 72:127–139. Gnandi, K., and Tobschall, H. J. 2003. Distribution patterns of rare-earth elements and uranium in tertiary sedimentary phosphorites of Hahotoé-Kpogamé, Togo. J. Afr. Earth Sci. 37:1–10. Haley, B. A.; Klinkhammer, G. P.; and McManus, J. 2004. Rare earth elements in pore waters of marine sediments. Geochim. Cosmochim. Acta 68:1265–1279. Hermina, M. 1990. The surroundings of Kharga, Dakhla and Farafra oases. In Said, R., ed. The geology of Egypt. Rotterdam, Balkema, p. 259–292. Howard, P. F., and Hough, M. J. 1979. On the geochemistry and origin of the D Tree, Wonarah, and Sherrin Creek phosphorite deposits of the Georgina Basin, northern Australia. Econ. Geol. 74:260–284. Ismael, I. S. 2002. Rare earth elements in Egyptian phosphorites. Chin. J. Geochem. 21:19–28. Issawi, B. 1972. Review of Upper Cretaceous-Lower Tertiary stratigraphy in central and southern Egypt. Am. Assoc. Petrol. Geol. Bull. 56:1448–1463. Jahnke, R. A. 1984. The synthesis and solubility of carbonate fluorapatite. Am. J. Sci. 284:58–78. Jarvis, I. 1992. Sedimentology, geochemistry and origin of phosphatic chalks: the Upper Cretaceous deposits of NW Europe. Sedimentology 39:55–97. Johannesson, K. H.; Palmore, C. D.; Fackrell, J.; Prouty, N. G.; Swarzenski, P. W.; Chevis, D. A.; Telfeyan, K.; White, C. D.; and Burdige, D. J. 2017. Rare earth element behavior during groundwater-seawater mixing along the Kona Coast of Hawaii. Geochim. Cosmochim. Acta 198:229–258. Kechiched, R.; Laouar, R.; Bruguier, O.; Salmi, S. L.; Zaimeche, Q. A.; and Foufou, A. 2016. Preliminary data of REE in Algerian phosphorites: a comparative study and paleo-redox insights. Proc. Eng. 138:19–29. Khan, K. F.; Dar, S. A.; and Khan, S. A. 2012. Geochemistry of phosphate bearing sedimentary rocks in parts of Sonrai Block, Lalitpur District, Uttar Pradesh, India. Chem. Erde 72:117–125. Klitzsch, E. 1978. Geological processing of southwest Egypt. Geol. Rundsch. 67:509–520. Lécuyer, C.; Reynard, B.; and Grandjean, P. 2004. Rare earth element evolution of Phanerozoic seawater recorded in biogenic apatites. Chem. Geol. 204:63–102. Lira, R., and Ripley, E. M. 1990. Fluid inclusion studied of the Rodeo de Los Molles REE and Th deposit, Las Chacras Batholith, central Argentina. Geochim. Cosmochim. Acta 54:663–671. Liu, P.; Liao, Y. C.; Yin, K. H.; and Ye, D. S. 2008. Hydrothermal sedimentary manganese deposits associated to volcanic activities—Permian manganese deposits in Guizhou. Geol. China 35:992–1006. Mariano, A. N. 1989a. Nature of economic mineralization in carbonatites and associated rocks. In Bell, K., ed. Carbonatites: genesis and evolution. London, Unwin-Hyman, p. 149–172. ———. 1989b. Economic geology of rare earth minerals. In Lipin, B. R., and McKay, G. A., eds. Geochemistry and mineralogy of rare earth elements. Rev. Mineral. 21. Washington DC, Mineralogical Society of America, p. 309–337. McArthur, J. M., and Walsh, J. N. 1985. Rare earth geochemistry of phosphorites. Chem. Geol. 47:191–220. McClellan, G. H., and Van Kauwenbergh, S. J. V. 1990. Mineralogy of sedimentary apatites. In Notholt, A. J. G., and Jarvis, I., eds. Phosphorite research and development. Geol. Soc. Lond. Spec. Publ. 52:23–31. Mgonde, F. 1994. Mobilization and redistribution of phosphate and rare earth elements in the weathering zone above Panda Hill carbonatite, SE Tanzania. MS thesis, Carleton University, Ottawa, Ontario, Canada, p. 27–33. Morad, S., and Felitsyn, S. 2001. Identification of primary Ce-anomaly signatures in fossil biogenic apatite: implication for the Cambrian oceanic anoxia and phosphogenesis. Sediment. Geol. 143:259–264. Nathan, Y. 1984. The mineralogy and geochemistry of phosphorites. In Nriagu, J. O., and Moore, P. B., eds. Phosphate minerals. Berlin, Springer, p. 276–281. Reynard, B.; Lécuyer, C.; and Grandjean, P. 1999. Crystal-chemical controls on rare-earth element concentrations in fossil biogenic apatites and implications for paleoenvironmental reconstructions. Chem. Geol. 155:233–241. Rona, P. A. 1984. Hydrothermal mineralization at seafloor spreading centers. Earth-Sci. Rev. 20:1–104. Ruttenberg, K. C., and Berner, R. A., 1993. Authigenic apatite formation and burial in sediments from non-upwelling, continental margin environments. Geochim. Cosmochim. Acta 57:991–1007. Sa, R.; Sun, X.; He, G.; Xu, L.; Pan, Q.; Liao, J.; Zhu, K.; and Deng, X. 2018. Enrichment of rare earth elements in siliceous sediments under slow deposition: a case study of the central North Pacific. Ore Geol. Rev. 29:12–23. Shatrov, V. A., and Voitsekhovskii, G. V. 2009. The use of lanthanides for the reconstruction of Phanerozoic and Proterozoic sedimentation environments exemplified by sections in the cover and basement of the East European platform. Geochem. Int. 47:758–776. Shields, G., and Stille, P. 2001. Diagenetic constraints on the use of cerium anomalies as paleoseawater redox proxies: an isotropic and REE study of Cambrian phosphorites. Chem. Geol. 175:29–48. Sholkovitz, E. R.; Elderfield, H.; Szymczak, R.; and Casey, K. A. 1999. Island weathering: river sources of rare earth elements to the Western Pacific Ocean. Mar. Chem. 68:39–57. Slansky, M. 1986. Geology of sedimentary phosphates, handbook. Cooper, P., trans. Oxford, North Oxford Academic. Sverjensky, D. A. 1984. Europium redox equilibria in aqueous-solution. Earth Planet. Sci. Lett. 67:70–78. Tantawy, A. A.; Keller, G.; Adatte, T.; Stinnesbeck, S. W.; and Kassab, A. 2001. Maastrichtian to Paleocene depositional environment of the Dakhla Formation, Western Desert, Egypt: sedimentology, mineralogy, and integrated micro- and macrofossil biostratigraphies. Cretaceous Res. 22:795–827. Uher, P.; Ondrejka, M.; Bačík, P.; Broska, I.; and Konečny, P. 2015. Britholite, monazite, REE carbonates, and calcite: products of hydrothermal alteration of allanite and apatite in A-type granite from Stupné, Western Carpathians, Slovakia. Lithos 236–237:212–225. Van Kauwenbergh, S. J.; Cathcart, J. B.; and McClellan, G. H. 1990. Mineralogy and alteration of the phosphate deposits of Florida: a detailed study of the mineralogy and chemistry of the phosphate deposits of Florida. U.S. Geol. Surv. Bull. 1914, 46 p. Wright, J.; Schrader, H.; and Holser, W. T. 1987. Paleoredox variations in ancient oceans recorded by rare earth elements in fossil apatite. Geochim. Cosmochim. Acta 51:631–644. Wu, C.; Yuan, Z. Z.; and Bai, G. 1996. Rare earth deposits in China. In Jones, A. P.; Wall, F.; and Williams, C. T., eds. Rare earth minerals—chemistry, origin and ore deposits. Mineral. Soc. Ser. 7:281–310. Youssef, M. I. 1957. Upper Cretaceous rocks in Kosseir area. Bull. Inst. Desert Egypt 7:35–53. Zabel, M., and Schulz, H. D. eds. 2006. Marine geochemistry. Berlin, Springer, p. 11–16.

By Galal El-Habaak; Mohamed Askalany and Mahmoud Abdel-Hakeem

Reported by Author; Author; Author