Geochemistry of Late Cretaceous phosphorites in Egypt: Implication for their genesis and diagenesis

Journal of African Earth Sciences 49 (2007) 12–28 www.elsevier.com/locate/jafrearsci

Geochemistry of Late Cretaceous phosphorites in Egypt: Implication for their genesis and diagenesis H.M. Baioumy *, R. Tada, M.H.M. Gharaie Department of Earth and Planetary Science, The University of Tokyo, Tokyo 113, Japan Received 23 March 2006; received in revised form 18 March 2007; accepted 8 May 2007 Available online 24 May 2007

Abstract Phosphorite deposits in Egypt, known as the Duwi Formation, are a part of the Middle East to North Africa phosphogenic province of Late Cretaceous to Paleogene age. Phosphatic grains in these deposites are classified into phosphatic mudclasts and phosphatic bioclasts. Phosphatic bioclasts are subdivided into fish bone fragments and shark tooth fragments. All phosphatic grains are composed of francolite. Chemical mapping of the phosphatic grains using Electron Probe Microanalysis (EPMA) indicated that the phosphatic mudclasts are homogeneous in their chemical composition and no concentric texture nor chemical zoning are observed. Some of the bone fragments show Fe and S zoning. No significant difference in chemical composition is observed between the phosphatic mudclasts and bioclasts. Acid-insoluble residues of the phosphorites show lower values of the Chemical Index of Alteration (CIA) compared to the associated rocks. Structural CO2 contents in the francolites range from 3.32% to 7.21% with an average of 5.3%. The d13CPDB values range from 4.04& to 8.7&, while the d18OPDB values range from 4.3& to 10.3&. The compositional homogeneity of the mudclasts, Fe and S zoning in some of the bone fragments and the difference in the Chemical Index of Alteration between the acid-insoluble residues of the phosphorites and the associated rocks suggest that the phosphatic grains in the Duwi Formation are derived from pre-existing authigenic phosphorites, which reworked and concentrated afterward. Negative d13C values of structural CO2 suggest that the CO2 was derived from degradation of organic matter. Low d18O values of structural CO2 can be attributed to the influence of meteoric water. Higher CO2, SO3 and F contents compared to the recent authigenic phosphorites and negative d13C and d18O values of structural CO2 indicate that diagenesis plays an important role in the modification of the chemical composition of phosphatic grains and that the studied apatite was francolitized during diagenesis.  2007 Elsevier Ltd. All rights reserved. Keywords: Egypt; Phosphorites; Phosphatic mudclasts; Phosphatic bioclasts; Francolite; Diagenesis

1. Introduction The geochemistry of phosphorites and their constituent mineral francolite have been widely studied owing to their economic importance (Jarvis et al., 1994) and the potential utility of their geochemistry to estimate paleo-marine chemistry (Donnelly et al., 1990). The chemical composition of carbonate fluorapatite is highly variable because *

Corresponding author. E-mail address: [email protected] (H.M. Baioumy).

1464-343X/$ - see front matter  2007 Elsevier Ltd. All rights reserved. doi:10.1016/j.jafrearsci.2007.05.003

its crystal structure allows a variety of substitutions (McClellan and Lehr, 1969). Variability in chemical composition of francolite may reflect difference in original composition, modification during diagenesis, or modification during weathering (Jarvis et al., 1994). Sedimentary francolite contains up to 8% of structurally incorporated carbonate (McArthur, 1978). This carbonate can be analyzed for its carbon and oxygen isotope composition, which are highly instructive on the condition of the francolite precipitation (Jarvis et al., 1994). The source of this carbon and oxygen in authigenic francolite is generally

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

considered as the pore-water at the time of its formation (Hudson, 1977). Because the amount of bicarbonate present in solution is small, its carbon isotope composition very sensitively reflects the degradation processes of organic matter during burial diagenesis. Consequently, authigenic francolite formed in anoxic sediments is expected to contain organically derived carbon released during bacterial degradation processes. The d13C of francolite should become light if it is precipitated in the sulfate-reduction zone whereas it should become heavy if it is precipitated in the fermentation zone (Hudson, 1977). Carbon isotopic data from francolite may be used to help constraining where and when phosphorite formation occurs (Glenn, 1990). If francolite precipitates directly from seawater at the sediment–water interface, then the carbon isotopic composition of incorporated CO2 should be characteristic of the bottom water. If, on the other hand, francolite precipitates within the sediments, the d13C values of lattice-bound CO2 should record d13C of the pore-water total dissolved carbon (Glenn et al., 1988). Degradation of marine organic matter typically contributes CO2 with a carbon isotopic composition of about 18& to 24& (cf. Anderson and Arthur, 1983) while the carbon isotopic composition of modern deep water total dissolved carbon is about 0.5& to +0.5& (cf. Kroopnick, 1985). The depletion in total dissolved carbon-13C in organic carbon-rich marine sediments takes place in suboxic to anoxic sulfate-reducing environments up to the point at which all dissolved sulfate has been used up by sulfate-reducing bacteria. Authigenic carbonatebearing mineral phases that precipitate within these zones may thus exhibit values ranging from about 0& for phases forming near the sediment–water interface, to values as low as 24& for phases forming under extensive sulfate-reduction. Concurrent with or following the sulfate-reduction processes, methane-producing bacteria use organic matter and/or total dissolved carbon in pore waters whereby extremely light (as light as 100&) biogenic methane is produced, leaving a residual total dissolved carbon enriched in 13C (Claypool and Kaplan, 1974). Therefore, below the zone of sulfate-reduction, methane production results in a progressive increase in total dissolved carbond13C from about 24& to +25&, depending upon the extent of methanogenesis. Authigenic mineral phases precipitated in this zone exhibits a range in d13C from very low to very high values (Glenn, 1990). The use of oxygen isotopes in marine carbonate minerals to estimate water paleotemperatures (Urey et al., 1951) and to determine diagenetic conditions (Hudson, 1977; Irwin et al., 1977) is well established. McArthur et al. (1980) interpreted d18O values of carbonate in apatite by using the calcite-water isotopic fractionation (Epstein et al., 1953), and reasonable temperatures were obtained for recent and young phosphorites. However, ancient phosphorites (from Jordan, South Africa, Phosphoria Formation) are consistently depleted in 18O.

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The Campanian-Maastrichtian phosphorite deposits in Egypt, named the Duwi Formation, form a portion of the Middle East to North Africa phosphogenic province of Late Cretaceous to Paleogene age (Notholt, 1985). This study aims to examine the mineralogical, chemical and isotopic compositions of the different phosphatic grains as well as temporal and spatial variations in the composition of phosphorites and associated rocks to assess their controlling factors. 2. Geological setting and studied localities The Duwi Formation overlies a fluvial shale sequence of the middle Campanian Qusseir Formation, and is overlain by the deeper marine shales and marls of the middle Maastrichtian Dakhla Formation. Thus, deposition of the Duwi Formation represents an initial stage of the late Cretaceous marine transgression in Egypt. The precise age of the Duwi Formation is poorly known, and generally considered as either late Campanian to early Maastrichtian based on paleontological evidences (Glenn and Arthur, 1990). According to Baioumy and Tada (2005), the Duwi Formation in the Red Sea, Nile Valley, and Abu-Tartur areas (Fig. 1) overlies non-marine, varicolored shale of the middle Campanian Qusseir Formation, and is comformably overlain by marine, laminated, gray, foraminefera-rich shale of the middle Masstrichtian Dakhla Formation. The Duwi Formation is subdivided into four members based on its lithology (Fig. 2). The lower member is composed of coarse phosphatic sandstone in the Abu-Tartur area whereas it is composed of quartzose sandstone and siliceous shale in the Nile Valley and Red Sea areas. The middle member is composed of soft, laminated, organicrich, black shale in the three areas. The upper member is composed of coarse glauconitic sandstone at Abu-Tartur area, phosphatic sandstone in the Nile Valley area, and phosphatic sandstone and oyster fragment-rich calcarenite in the Red Sea area, respectively. The uppermost member is composed of hard, massive grayish brown to gray shale in the three areas. Individual phosphorite beds in the Duwi Formation range in thickness from a few millimeters to tens of centimeters. Thicker phosphorite beds are formed by amalgamation of thinner individual beds. The thickest accumulation of minable phosphorites occurs in the lower member in Abu-Tartur area where the phosphorite beds locally amalgamate to form a single seam averaging approximately 12 m thick. One common feature of nearly all Duwi phosphorites is extensive bioturbation. As a result, most of the phosphatic beds appear massive and internally structureless. 3. Samples and methods Thin sections of phosphorites and the associated rocks were prepared and observed under the petrographic microscope to examine their grain composition and texture. A total of 42 phosphatic grains from seven polished thin

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Fig. 1. Geological map of Egypt with the localities of the studied areas (modified from Spanderashvilli and Mansour, 1970).

sections of the phosphorites that represent two different phosphorite horizons from the three areas were analyzed by electron probe microanalyser (EPMA) for their major (SiO2, TiO2, Al2O3, FeO, MnO, MgO, CaO, K2O, Na2O, SO3, P2O5, F and Cl) and trace (NiO, Cr2O3 and V2O3) element composition. Samples for EPMA analysis were selected based on the microscopic observations to represent all types of phosphatic grains. Analyses were conduced using a JEOL, JXA-8900L Electron Probe Microanalyzer at the Department of Earth and Planetary Science, University of Tokyo, which was operated at accelerating voltage of 15 kV, beam diameter of 5 lm, and beam current of 1.2 · 108 mA. Five spots across each grain have been analyzed to check the presence of compositional zoning. Under the same operating conditions, eleven phosphatic grains have been mapped for the major elements (SiO2, Al2O3, FeO, MnO, MgO, CaO, K2O, Na2O, SO3, P2O5, and F), to visually examine compositional variation within grains.

Acid-insoluble residues of 17 samples of phosphorites and 35 samples of the associated rocks were analyzed for the major elements (SiO2, TiO2, Al2O3, FeO, MnO, MgO, CaO, K2O, Na2O, and P2O5) by X-ray fluorescence using a Philips X-ray spectrometer equipped with a Rh tube in the Department of Earth and Planetary Science, University of Tokyo. Tube voltage and current were 40 kV and 60 mA, respectively. Chemical Index of Alteration (CIA) according to Nesbitt and Young (1982) is calculated for insoluble residues of phosphorites, shales, and calcarenite in order to evaluate the degree of chemical weathering of detrital materials in these rocks, which is considered to reflect climate in the detrital source area. CIA is calculated by the following formula (Nesbitt and Young, 1982): CIA = [Al2O3/(Al2O3 + CaO + Na2O + K2O)] · 100. Eleven of physically separated phosphatic grains were analyzed by X-ray powder diffractometry (XRD) to

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

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Fig. 2. Correlation of columnar sections of the Duwi Formation and equivalent phosphate-bearing formations in the studied localities (from Baioumy and Tada, 2005).

examine their mineral composition. Powdered samples were mounted on a glass holder and X-rayed from 2 to 80 2h at 2 2h/min. using a MAC Science MXP-3 X-ray diffractometer equipped with a CuKa tube at the Department of Earth and Planetary Science, University of Tokyo. Tube voltage and current are 40 kV and 20 mA, respectively. The replacement of the tetrahedral PO4 by a CO3 anion changes the lattice parameter of the apatite unit cell (LeGeros et al., 1967). This relationship was used by Gulbrandsen (1970) and Scuffert et al. (1990) to estimate the CO2 content of apatite. With the increase in substitution of phosphate by carbonate, the (4 1 0) plane shifts to a higher 2h angle and the (0 0 4) plane shifts to a lower 2h angle; i.e. the angular distance between the two peaks

decreased. This angular distance between the two peaks, which are approximately at 51.6 and 53.1 2h respectively, is used as a measure of the CO2 content in the separated phosphatic mudclasts and phosphatic bioclasts according to the equation of Scuffert et al. (1990): Y ¼ 10:643X 2  52:512X þ 56:986; where Y = CO2% and X = D2h(0 0 4)–(4 1 0) A total of 24 phosphorite samples were analyzed for their d13C and d18O compositions. To extract CO2 for isotopic analysis, 20 mg of powdered sample were reacted with 2.5 ml high concentrated phosphoric acid in a reaction tube under vacuum condition at 25 C for about 12 h. CO2 was recovered and analyzed by a MAT252 mass spectrometer

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at the Earth and Planetary Science, the University of Tokyo. Carbon and oxygen isotopes data are reported in & notation relative to the PDB standard. 4. Results 4.1. Petrography of the phosphorites Phosphorites are composed of phosphatic grains, nonphosphatic grains, cements, and pores. Phosphatic grains are composed of phosphatic mudclasts and bioclasts. Phosphatic mudclasts are generally well-rounded to subrounded, and spherical in shape. Their size ranges from 40 to 100 lm in diameter, and colorless transparent through yellow to brown translucent. Black opaque varieties are confined to fresh samples from the Abu-Tartur mine. Although phosphatic mudclasts are spherical in shape, no concentric texture nor any other internal structure is observed (Fig. 3). The percentage of phosphatic mudclasts in the phosphatic grains range from 45 to 67 vol% with an average of 51 vol% in the Duwi Formation (Baioumy and Tada, 2005). Phosphatic bioclasts are composed of fragments of fish bone and shark tooth. Fish bone fragments are angular to subangular and prismatic or irregular in shape (Fig. 4). Their size ranges from 120 to 180 lm. Generally, bone fragments are colorless and transparent. However, some of bone fragments show various shades of color ranging from yellow to gray. Although the majority of bone fragments is internally structureless, some of the bone fragments in the phosphorite of the lower member in Abu-Tartur area show color zoning. In such bone fragments,

disseminated pyrite occurs in the core, which is surrounded by a yellowish oxidized rim (Fig. 5). Bone fragments are generally anisotropic, but some show low birefringence and lamellar twinning. Elongated and prismatic bone fragments occasionally show extinction parallel to the elongated axis. The percentage of bone fragments in the phosphatic grains of the Duwi phosphorites range from 36 to 54 vol% with an average of 43.7 vol% (Baioumy and Tada, 2005). Tooth fragments (Fig. 6) are less abundant compared to bone fragments. The color of tooth fragments ranges from yellow to brown translucent and generally anisotropic. Size of tooth fragments ranges from 150 to 200 lm. Larger tooth fragments are generally rounded, whereas the smaller are irregular in shape with angular edges probably due to fragmentation. Generally, tooth fragments are larger than other types of phosphatic grains. The percentage of tooth fragments in the total phosphatic grains of the Duwi phosphorites range from 0.25 to 2 vol% with an average of 1.3 vol% (Baioumy and Tada, 2005). Non-phosphatic grains are mostly composed of subangular to subrounded, mono- and polycrystalline, colorless to pale gray, detrital quartz and rounded, green to greenish gray glauconite pellets. Cements are composed of chalcedony, calcite, dolomite, pyrite, gypsum, anhydrite, and iron oxides. 4.2. Mineralogy of phosphatic grains Mineralogical analysis by XRD identified seven mineral phases in the Duwi phosphorites including francolite, calcite, dolomite, quartz, gypsum, anhydrite, and pyrite.

Fig. 3. A thin section photomicrograph of a well-rounded and internally structureless phosphatic mudclast from the lower member of the Duwi Formation in Abu-Tartur area.

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Fig. 4. A thin section photomicrograph of a fish bone fragment (arrows) from the lower member of the Duwi Formation in the Nile Valley area.

Fig. 5. A thin section photomicrograph of a fish bone fragment with pyritized core and oxidized rim from the lower member of the Duwi Formation in the Abu-Tartur area.

XRD analysis revealed that both phosphatic mudclasts and phosphatic bioclasts are composed of francolite. The X-ray diffraction patterns of both grains show relatively broad and poorly resolved peaks of francolite reflecting low crystalinity of the francolite. The peak position of francolite in the Duwi Formation shows slight shifts from the position of typical francolite, which suggests changes in the cell parameters as a result of isomorphous substitutions. 4.3. Chemistry of phosphatic grains Phosphatic grains from different localities and horizons by EPMA to examine possible differences in chemical com-

position between different types of phosphatic grains from different horizons and localities. Average composition and standard deviations of the three major grain types from the two phosphatic horizons at the studied localities are summarized in Table 1. Chemical composition of the three grain types from the same horizon in the same area is similar to one another within the analytical error, whereas slight difference is observed between different horizons and localities. For example, phosphatic grains from the lower member in the Red Sea and Nile Valley areas show higher SO3 content compared to those from the upper member. At Abu-Tartur area, the phosphatic grains from the upper member have

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Fig. 6. A thin section photomicrograph of a shark tooth from the upper member of the Duwi Formation in Abu-Tartur area.

lower Al2O3, MnO, and K2O contents compared to those in the lower member as well as those in other areas. Phosphatic grains from both the upper and lower members in Abu-Tartur area generally show higher Na2O contents compared to those in the Red Sea and Nile Valley areas. The result of chemical mapping revealed that Ca, P, F, S, Na, Mn, and Mg show homogeneous distributions, suggesting that they are held as structural components (Fig. 7). On the other hand, Si, K, Fe, and Al are concentrated in small spots, suggesting that they occur as microinclusions (Fig. 10). Mapping analysis also shows that the phosphatic mudclasts are generally homogeneous and do not show any compositional zoning (Fig. 8). Some of bone fragments show false zonation of FeO and SO3 due to replacement followed by leaching of pyrite from the grain rims (Fig. 9). 4.4. Chemistry of acid-insoluble residues of phosphorites and associated rocks Acid-insoluble residues of the phosphorites and their associated rocks were analyzed for the major elements by XRF to examine the potential difference in detrital source between the phosphorites and the associated rocks. Table 2 shows the analytical results. Acid-insoluble residues of phosphorites tend to show higher SiO2 contents and lower Al2O3, TiO2, MgO, and K2O contents compared to those of other rock types. This probably reflects presence of chalcedony cement in phosphorites that dilute other elements. By assuming that Al in the residues is exclusively derived from detrital aluminosilicate sources (Murray et al., 1992), the relation between Al2O3 and other elements are examined to check the possible difference in provenance. Positive

correlations are observed between Al2O3 and TiO2 (Fig. 10) in the acid-insoluble residues of phosphorites and their associated shales and carbonates, while positive correlation between Al2O3 and K2O (Fig. 11) in the acid-insoluble residues of phosphorites only. This correlation is weak in the associated shales and carbonates. The Al-normalized ratios of these elements in the acid-insoluble residues of phosphorites are different from those of the associated rocks. For example, residues of most of phosphorites has TiO2/Al2O3 larger than 0.09, which is larger than the ratio for insoluble residues in other rock types that generally have ratios of 0.04–0.07 (Fig. 10). Many of phosphorite residues have K2O/Al2O3 larger than 0.1, which is larger than other rock types except glauconitic samples (Fig. 11). The differences in Ti/Al and K/Al ratios between the acid insoluble residue of phosphorites and the associated rocks can be attributed to the difference in the source of detrital material between phosphatic grains and associated rocks in the Duwi Formation. Fig. 12 shows ranges and average values of CIA for the acid-insoluble residues of different rock types. Acid-insoluble residues of the phosphorites show lower values of (CIA) compared to the associated shales. 4.5. Structural CO2 and its carbon and oxygen isotopic compositions The structural CO2 contents of francolite in the studied samples range between 3.3% and 7.2% with an average of 4.9% (Table 3). The d13C values range from 4& to 8.7& with an average of 5.9&, while the d18O values range from 4.4& to 10.3& with an average of 7.4& (Table 3). Slight but clear differences can be observed in

Table 1 Mean values (wt.%) and standard deviation for the three different types of phosphatic grains from the two members in the three different localities Elements

SiO2

Al2O3 FeO MnO MgO CaO K2O Na2O Ni2O3 Cr2O3 V2O3 SO3 P2O5 CO2 F Cl Total

Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave Ave STDV Ave STDV Ave STDV

Red Sea, lower member

Nile Valley, upper member

Nile Valley, lower member

Bone

Tooth

Structureless grain

Bone

Tooth

Structureless grain

Bone

Tooth

Structureless grains

0.09 0.04 0.03 0.03 0.02 0.01 0.35 0.53 0.01 0.02 0.10 0.01 53.41 0.45 0.01 0.01 0.24 0.08 0.06 0.04 0.01 0.01 0.01 0.01 0.59 0.51 35.58 0.76 4.92 4.56 0.47 0.01 0.01 100.00 0.75

0.07 0.03 0.04 0.02 0.02 0.02 0.65 0.24 0.02 0.03 0.16 0.02 53.11 0.24 0.01 0.01 0.45 0.09 0.03 0.03 0.02 0.03 0.01 0.01 1.08 0.63 35. 25 0.16 4.98 4.11 0.66 0.01 0.00 100.00 0.83

0.03 0.03 0.00 0.00 0.01 0.02 0.18 0.04 0.01 0.02 0.15 0.02 53.26 0.36 0.02 0.01 0.49 0.09 0.03 0.03 0.02 0.01 0.01 0.02 1.37 0.07 35.11 0.81 4.95 4.347 0.29 0.02 0.02 100.00 1.39

0.11 0.03 0.01 0.01 0.01 0.01 0.32 0.13 0.01 0.02 0.17 0.04 53.34 0.50 0.03 0.01 0.53 0.03 0.03 0.03 0.04 0.07 0.02 0.01 1.90 0.14 33.10 0.65 6.12 4.27 0.37 0.03 0.02 100.00 1.30

0.08 0.03 0.02 0.03 0.01 0.01 0.23 0.01 0.01 0.00 0.11 0.02 52.29 0.28 0.03 0.01 0.54 0.05 0.00 0.01 0.00 0.01 0.00 0.01 2.10 0.12 33.99 0.47 6.02 4.54 0.52 0.09 0.03 100.00 0.80

0.09 0.03 0.00 0.00 0.00 0.00 0.13 0.03 0.03 0.00 0.17 0.02 53.12 0.43 0.02 0.05 0.46 0.10 0.03 0.06 0.01 0.02 0.02 0.08 1.16 0.34 35.24 0.90 5.10 4.41 0.43 0.02 0.00 100.00 0.45

0.06 0.02 0.00 0.00 0.01 0.01 0.06 0.03 0.00 0.00 0.13 0.03 52.45 0.32 0.03 0.01 0.41 0.08 0.03 0.05 0.01 0.01 0.01 0.01 1.19 0.04 35.62 0.94 5.23 4.70 0.65 0.05 0.03 100.00 0.33

0.03 0.01 0.01 0.02 0.02 0.02 0.20 0.23 0.09 0.06 0.19 0.03 53.25 0.04 0.05 0.01 0.46 0.06 0.01 0.02 0.03 0.03 0.00 0.00 1.10 0.05 35.18 0.91 5.21 4.15 0.37 0.04 0.02 100.00 0.36

0.04 0.00 0.01 0.00 0.01 0.01 0.16 0.00 0.02 0.04 0.13 0.00 52.19 0.00 0.07 0.01 0.75 0.07 0.05 0.00 0.01 0.00 0.02 0.00 2.20 0.05 34.60 0.00 5.73 4.02 0.00 0.04 0.00 100.00 0.00

0.15 0.00 0.00 0.00 0.07 0.02 1.22 0.00 0.20 0.03 0.24 0.00 51.95 0.00 0.03 0.02 0.83 0.06 0.10 0.00 0.01 0.00 0.00 0.00 1.69 0.06 33.81 0.00 5.73 3.98 0.00 0.09 0.00 100.00 0.00

Bone

Tooth

0.01 0.08 0.02 0.00 0.01 0.04 0.01 0.01 0.01 0.01 0.01 0.00 0.39 0.28 0.44 0.00 0.02 0.08 0.02 0.03 0.13 0.14 0.02 0.00 52.65 52.03 0.56 0.00 0.04 0.03 0.02 0.01 0.65 0.40 0.05 0.03 0.03 0.00 0.04 0.00 0.00 0.04 0.00 0.00 0.03 0.00 0.03 0.00 1.29 1.9 0.15 0.24 34.65 34.78 0.40 0.00 5.77 6.05 4.31 4.13 0.72 0.49 0.08 0.09 0.02 0.00 100.00 100.00 0.76 0.00 (continued on next page)

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TiO2

Red Sea, upper member Structureless grains

19

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Table 1 (continued) Elements

SiO2 TiO2

FeO MnO MgO CaO K2O Na2O Ni2O3 Cr2O3 V2O3 SO3 P2O5 CO2 F Cl Total

Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave STDV Ave Ave STDV Ave STDV Ave STDV

Abu-Tartur, lower member

Abu-Tartur, fresh sample (lower member)

Structureless grains

Bone

Tooth

Structureless grain

Bone

Tooth

Structureless grain

Bone

Tooth

0.05 0.03 0.02 0.03 0.02 0.02 0.39 0.35 0.05 0.04 0.12 0.03 52.72 0.57 0.04 0.03 1.02 0.06 0.01 0.02 0.01 0.01 0.01 0.02 2.16 0.07 35.87 0.37 3.62 3.80 0.82 0.03 0.01 100.00 1.07

0.03 0.04 0.01 0.01 0.01 0.01 0.79 0.38 0.08 0.05 0.25 0.07 51.93 0.55 0.01 0.00 1.08 0.08 0.03 0.06 0.01 0.03 0.02 0.02 2.33 0.23 35.81 0.45 3.60 4.00 0.31 0.05 0.01 100.00 0.64

0.08 0.04 0.00 0.01 0.03 0.02 1.62 0.40 0.09 0.06 0.14 0.03 51.54 1.11 0.07 0.04 0.91 0.11 0.01 0.02 0.01 0.01 0.07 0.01 2.29 0.35 35.52 0.82 3.62 3.90 0.95 0.05 0.03 100.00 0.98

0.06 0.00 0.02 0.00 0.83 0.00 0.98 0.04 0.25 0.03 0.30 0.00 51.24 0.00 0.06 0.02 0.94 0.21 0.00 0.00 0.01 0.00 0.02 0.00 1.90 0.35 34.82 0.00 4.34 4.22 0.00 0.03 0.00 100.00 0.00

0.26 0.31 0.00 0.02 0.77 0.10 0.91 0.32 0.28 0.07 0.31 0.03 51.59 0.69 0.06 0.01 0.92 0.10 0.00 0.04 0.04 0.02 0.01 0.01 1.92 0.14 34.34 0.75 4.38 4.10 0.53 0.04 0.02 100.00 1.07

0.03 0.02 0.01 0.02 1.02 0.04 0.65 0.07 0.07 0.02 0.24 0.03 51.16 0.68 0.06 0.01 0.77 0.04 0.03 0.03 0.03 0.02 0.01 0.01 1.80 0.12 35.69 0.84 4.37 4.00 0.51 0.02 0.01 100.00 1.31

0.05 0.06 0.02 0.02 0.90 0.01 1.70 0.19 0.20 0.19 0.30 0.05 50.93 0.07 0.04 0.01 0.89 0.15 0.01 0.02 0.00 0.00 0.01 0.01 4.20 0.30 34.19 0.30 4.54 3.88 0.55 0.04 0.02 100.00 0.93

0.01 0.02 0.01 0.01 0.01 0.01 0.68 0.70 0.07 0.01 0.15 0.03 51.86 0.65 0.03 0.01 0.89 0.17 0.04 0.06 0.03 0.04 0.01 0.01 2.96 0.31 34.49 1.34 4.52 4.24 0.52 0.04 0.01 100.00 1.22

0.01 0.01 0.01 0.01 0.01 0.01 1.53 0.25 0.05 0.02 0.18 0.02 51.08 0.25 0.04 0.01 1.03 0.05 0.01 0.02 0.02 0.01 0.01 0.00 3.66 0.46 33.75 0.15 4.50 4.09 0.43 0.03 0.01 100.00 0.90

Ave: average. STDV: standard deviation.

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Al2O3

Abu-Tartur, upper member

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Fig. 7. Chemical composition maps of a phosphatic mudclast from the lower member in Abu-Tartur area measured by EPMA. Ca, P, F, S, Na, Mg, and Mn are homogeneously distributed within the grain, suggesting their occurrence as lattice held element in the francolite structure.

the CO2%, d18O, and d13C values among the studied localities (Fig. 13). Samples from the Nile Valley area have higher CO2% and lower d13C values compared to those from the Red Sea and Abu-Tartur areas, whereas, samples from the Red Sea area have lower CO2% and higher d13C and lower d18O values compared to those from the Nile Valley and Abu-Tartur areas. Abu-Tartur samples have

intermediate values. Scatter diagram of CO2 contents and d13C is shown in Fig. 14. A linear relationship between d13C and CO2% is observed with different trends among different localities and horizons ranging from very strong correlation (r2 = 0.9) in case of Red Sea lower member phosphorites to very week correlation (r2 = 0.2) in case of Nile Valley lower member phosphorites.

22

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

Fig. 8. Chemical composition maps of a phosphatic mudclast from the lower member in Abu-Tartur area measured by EPMA. Al, Si, K, and Fe are found only as patches within the grain, suggesting their occurrence as microinclusions.

Fig. 9. Chemical composition maps of a phosphatic bioclast (fish bone fragment) from the lower member in Abu-Tartur area measured by EPMA. It shows false zonation of FeO (a) and SO3 (b) due to replacement followed by leaching of pyrite from the grain rims.

5. Discussion 5.1. Genesis of the phosphorites Late Cretaceous phosphorites in Egypt are made of two basic grain types; phosphatic mudclasts and bioclasts.

Chemical mapping of the phosphatic mudclasts indicated no concentric structure that characterizes pellets of in situ formation and phosphatic mudclasts are homogeneous in their chemical composition. It also indicated the presence of Fe and S zoning of some bone fragments. These observations in addition to the occurrence of detrital grains in

Table 2 Results of XRF analysis (wt.%) of acid-insoluble residue (AIR) of the phosphorites and associated rocks in the Duwi Formation Lithology

Samples

A.I.R.

L.O.I.

SiO2

TiO2

Al2O3

Fe2O3

MnO

MgO

CaO

Na2O

K2O

P2O5

Total

Al2O3/TiO2

CIA

Red Sea

Phosphorites

Rph6 Rph8 Rph9 Rph14 Rph22 Nph37 Nph39 Nph42 Nph43 Nph49 Nph52 Nph53 Nph59 Nph62 Aph75 Aph82 Aph85

3.69 4.24 30.48 8.76 22.67 9.64 7.94 33.76 33.02 12.57 39.58 23.91 26.20 41.59 39.92 22.30 40.51

2.84 3.61 2.76 6.91 2.56 4.71 3.99 4.47 2.92 2.40 2.89 16.72 2.13 1.80 23.06 13.26 17.74

87.78 86.93 85.76 85.04 94.14 84.47 83.53 87.92 90.23 95.16 91.96 68.04 96.24 97.10 59.39 69.15 66.99

0.04 0.08 0.11 0.02 0.08 0.30 0.10 0.17 0.07 0.02 0.02 0.02 0.01 0.03 0.38 0.51 0.51

0.28 1.04 0.18 0.31 1.37 2.23 0.38 1.39 0.40 0.16 0.49 2.08 0.19 0.69 1.13 4.96 3.48

7.26 5.72 10.64 0.64 1.33 6.89 11.08 4.74 5.73 2.15 1.53 9.64 1.30 0.22 10.46 9.56 9.41

0.03 0.01 0.07 0.01 0.01 0.01 0.03 0.02 0.03 0.01 0.01 0.00 0.01 0.01 0.03 0.02 0.02

0.15 0.18 0.17 0.12 0.19 0.22 0.14 0.17 0.14 0.10 0.03 0.03 0.10 0.09 0.19 0.85 0.39

1.35 1.71 0.09 6.51 0.07 1.95 0.17 0.11 0.07 0.09 2.55 2.18 0.20 0.05 4.17 0.86 0.79

0.16 0.47 0.44 0.13 0.19 0.34 0.15 0.30 0.17 0.12 0.06 0.05 0.12 0.69 0.11 0.18 0.00

0.03 0.07 0.05 0.03 0.20 0.33 0.03 0.18 0.05 0.02 0.03 0.02 0.02 0.07 0.20 0.45 0.37

0.24 0.28 0.14 0.52 0.04 0.25 0.18 0.22 0.16 0.07 0.24 0.21 0.12 0.02 0.29 0.20 0.13

100.17 100.11 100.41 100.23 100.17 101.70 99.79 99.70 99.95 100.30 99.79 98.99 100.45 100.78 99.40 100.00 99.82

6.52 13.52 1.69 17.68 18.30 7.47 3.95 8.34 5.55 9.22 26.42 131.68 13.64 22.71 3.02 9.74 6.90

15.34 31.60 23.80 4.48 75.03 73.98 51.43 70.08 58.51 40.99 15.61 47.97 34.92 46.04 20.20 76.83 74.98

Shales

RDsh2 RDsh3 RDsh4 RDsh5 NDsh2 NDsh4 NDsh7 NDsh12 NDsh14

97.47 93.90 88.52 52.21 95.94 94.87 92.07 95.22 94.20

19.80 16.67 15.28 15.66 15.01 19.84 16.08 10.33 15.40

57.80 57.93 60.35 57.68 59.06 53.31 59.32 71.37 63.99

0.69 1.00 0.79 0.99 1.01 0.95 0.86 0.61 0.62

13.94 16.76 13.22 17.16 17.57 17.20 14.01 10.76 12.77

4.55 3.45 5.96 3.77 3.70 4.70 6.14 4.41 4.27

0.01 0.01 0.01 0.01 0.01 0.01 0.09 0.01 0.01

1.93 1.82 1.60 1.47 1.51 1.67 1.42 1.25 1.34

0.00 0.00 0.01 0.00 0.00 0.00 0.01 0.00 0.00

0.29 0.07 0.15 0.35 0.00 0.01 0.00 0.00 0.00

0.65 0.74 1.17 0.80 0.01 0.01 0.01 0.01 0.01

0.08 0.61 0.22 0.13 0.13 0.05 0.19 0.07 0.10

99.74 99.05 98.76 98.02 98.03 97.74 98.12 98.82 98.49

20.35 16.80 16.65 17.31 17.31 18.12 16.22 17.67 20.65

93.65 95.38 90.86 93.71 99.91 99.89 99.85 99.90 99.89

Abu-Tartur

Shales

ABsh1 ABsh2 ABsh3 ABsh4 ABsh5 ABsh6 ABsh7 ABsh8 ABsh9 ABsh10

86.03 88.80 85.78 89.30 87.28 94.54 89.00 91.83 83.79 94.86

19.65 25.18 14.80 18.68 19.91 22.01 20.05 23.37 21.05 24.90

54.79 50.08 57.52 52.25 53.71 51.55 53.72 47.76 51.34 53.01

0.81 0.91 0.85 0.87 0.80 0.77 0.92 0.79 0.73 0.77

16.54 13.33 17.88 16.97 16.40 16.31 16.63 17.07 15.86 15.54

4.74 5.74 5.19 6.21 5.26 4.94 5.67 5.69 6.04 3.99

0.02 0.01 0.02 0.02 0.03 0.02 0.01 0.01 0.02 0.01

1.92 1.80 2.09 2.23 2.00 1.94 2.08 2.14 2.23 1.72

0.54 0.18 0.46 0.35 0.50 0.36 0.27 0.35 0.77 0.25

0.04 0.04 0.07 0.35 0.08 0.19 0.02 0.18 0.05 0.00

1.23 0.91 1.34 1.02 1.20 1.18 0.79 0.72 1.09 0.76

0.28 0.05 0.21 0.11 0.15 0.12 0.05 0.07 0.20 0.05

100.56 98.23 100.44 99.04 100.04 99.40 100.21 98.14 99.37 101.00

20.42 14.61 20.92 19.50 20.46 21.19 18.08 21.72 21.75 20.26

90.14 92.21 90.53 90.82 90.21 90.36 93.93 93.16 89.29 93.84

Red Sea

Carbonate

RM1(D) RM2(D) RM3(L) RDsh1

15.7 11.0 12.5 71.1

26.4 14.0 13.8 18.6

60.7 61.4 67.8 50.6

0.3 0.5 0.4 0.7

6.8 9.9 10.0 15.5

3.2 8.9 4.6 8.0

0.1 0.0 0.0 0.0

1.1 1.6 1.5 2.8

0.4 0.0 0.0 1.3

0.0 0.0 0.1 0.1

0.3 1.3 0.4 1.1

0.4 0.2 0.0 0.1

99.8 97.8 98.6 98.9

23.3 18.6 23.9 22.9 (continued on

89.5 88.1 95.1 86.1 next page)

Nile Valley

Abu-Tartur

Red Sea

Nile Valley

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

Area

23

17.0 16.2 14.5 14.6 17.3 25.6 25.4 24.9 20.6 12.2 12.1

Al2O3/TiO2 Total P2O5

0.1 0.1 0.1 0.1 0.0 0.1 0.5 0.2 1.7 1.0 1.1

K2O

0.0 0.0 0.0 0.0 0.3 3.4 4.1 3.1 3.4 1.4 1.5

Na2O

0.0 0.0 0.0 0.0 0.1 0.1 0.1 0.0 0.3 0.0 0.0

CaO

0.0 0.0 0.0 0.0 0.2 0.2 0.4 0.1 0.1 1.3 1.4 1.2 0.7 1.0 1.0 0.8 2.2 1.8 2.5 1.6 1.5 1.5 Glauconite Abu-Tartur

MgO MnO

0.1 0.1 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 6.1 4.9 4.8 4.8 3.7 18.3 27.5 12.8 16.4 6.8 7.1

Fe2O3 Al2O3

11.6 5.5 6.9 7.1 7.8 7.6 4.7 10.1 5.1 8.5 8.9 0.7 0.3 0.5 0.5 0.5 0.3 0.2 0.4 0.2 0.7 0.7

TiO2 SiO2

61.1 79.8 75.5 76.8 76.9 58.8 50.7 63.2 63.8 61.6 64.5 16.4 8.4 10.0 9.9 9.4 10.5 10.5 8.4 8.4 16.6 13.6

L.O.I.

mudclasts and fragments of oval-shaped mudclasts suggest that the phosphatic mudclasts in the Duwi Formation are not formed in situ by chemical precipitation of their constituents from solutions but derived along with the phosphatic bioclasts from pre-existing authigenic phosphorites formed outside the depositional basin and then reworked and concentrated in the depositional basin. During reworking, the margins of pyritic bone fragments were oxidized while cores still preserve their pyrite. The difference in the values of Chemical Index of Alteration (CIA) between the acid-insoluble residues of the phosphorites and the associated rocks suggests less intense weathering during the formation of phosphatic grains compared with that for associated rocks in the Duwi

66.9 77.0 43.0 69.1 86.6 91.2 81.6 78.5 83.8 67.6 73.3

A.I.R.

Fig. 10. Al2O3 vs TiO2 relation of acid-insoluble residue of phosphorites and their associated rocks. The trend for phosphorites is different from that for associated rocks possibly reflecting change in source of detrital influx.

NM4(M) NM5(M) NM6(M) NM7(M) NM8(M) AGl1 AGl2 AGl3 AGl4 AGl6 AGl7

Samples Lithology Area

Nile Valley

Table 2 (continued)

97.4 99.8 98.7 100.2 99.7 101.5 100.6 100.9 100.8 99.4 100.3

CIA

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

99.9 99.9 99.9 99.9 92.9 67.8 50.3 75.7 58.0 75.5 75.5

24

Fig. 11. Al2O3 vs K2O relation of acid-insoluble residue of phosphorites and their associated rocks. The trend for phosphorites is different from that for associated rocks possibly reflecting change in source of detrital influx.

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

Fig. 12. Variation in the values of Chemical Index of Alteration (CIA) for acid-insoluble residues of different rock types in the Duwi Formation. Phosphorites show lower CIA values compared to other rock types.

Formation. This could reflect different climate at the time of phosphorite formation either due to the difference in timing or place of formation, supporting the proposed formational mechanism of the Duwi phosphorites that the phosphatic grains are derived from pre-existing authigenic phosphorites formed and then reworked. Baioumy and Tada (2005) suggested that the pre-existing authigenic phosphatic mud was formed under an upwelling area of Tethyan margin during Early Campanian high stands and then reworked to a continental margin environment during the subsequent transgression based on geological and sedimentological evidences of these phosphorites.

Table 3 Structural CO2 (wt.%), d13C, d18O (& ) of the Duwi phosphorites Locality

Samples

CO2

d13C

Red Sea

PR4 PR8 PR9 PR12 PR14 PR15

6.58 6.58 7.21 6.23 6.21 6.71

7.91 7.74 8.71 7.71 5.25 6.18

6.4 6.38 5.9 7.91 5.96 7.25

Nile Valley

PN37 PN40 PN48 PN53 PN54 PN58 PN60 PN61 PN63

4 4.25 5.21 3.32 6.7 3.78 3.68 3.72 4.87

5.2 4.75 4.83 4.04 5.28 4.51 4.46 4.65 5.55

9.9 9.5 9.41 10.3 8.8 9.15 9.44 10.3 9.9

PA79 PA85 PA100 PAW2 PAW1 PAB47 PAB43 PAB41 PAB40

5.78 5.95 5.87 5.25 4.78 4.88 4.44 5.33 5.3

6.31 6.56 6.45 6.1 5.67 6.1 5.94 6.21 6.1

4.37 6.33 6.36 5.94 6.27 4.93 4.99 6.23 4.84

Abu-Tartur

d18O

25

Fig. 13. Histogram shows a comparison of the structural CO2%, d13C and d18O average values of the studied phosphorites among the studied localities.

As is suggested from the chemical mapping analysis, Mg is considered as a structural element held in francolite lattice. Consequently, variation in its content can be attributed to the difference in formation condition of francolite in the phosphorites. According to Baturin and Bezrukov (1979), authigenic phosphatic nodules of Peru and Chile margins were formed under upwelling conditions similar to those of the formational conditions of the Duwi phosphatic mudclasts proposed by Baioumy and Tada (2005). The lower MgO content in the francolite in the phosphorites of the Duwi Formation (average = 0.18%) compared the Chile and Peru margins authigenic phosphorites proposed by Baturin and Bezrukov (1979) (MgO% = 1.5% and 2.3% respectively) can be attributed to the formation of these phosphatic grains in low MgO sea water. According to Sandberg (1983), Mg content was generally low during Cretaceous. Soudry (1992) reported that the authigenic phosphorites from Campanian Mishash Formation in Israel are intercalated with authigenic montmorillonite, which might have reduced Mg ions from the pore-water during early diagenesis. The low Mg contents in the Duwi phosphorites could be used as evidence that supports the interpretation of Baioumy and Tada (2005) that the phosphatic mudclasts in the Duwi phosphorites were derived from authigenic phosphorites such as the authigenic phosphorites from Campanian Mishash Formation. 5.2. Diagenesis of the phosphorites McArthur (1985), Jarvis et al. (1994) and Kolodny and Luz (1992) proposed several controlling factors for the variation in francolite composition. Namely, mechanism of formation, formation environment, diagenetic alterations, metamorphism, and weathering. According to Baturin and Bezrukov (1979), the CO2 contents of authigenic phosphatic nodules of Peru and Chile margins range between 3.02% and 3.57% and F contents between 1.6% and 2.55%, which are lower than those

26

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

Fig. 14. Scatter diagram of CO2 contents and d13C. A linear relationship between d13C values and CO2% is observed.

in the Egyptian phosphorites (4.8% and 4.25%, respectively, Table 3). Gulbrandsen (1970) attributed regional variations in CO2 content of apatite in the Phosphoria Formation to temperature variations affecting CO2 saturation. Romankevich and Baturin (1972) suggested that the carbonate is supplied from the degradation of organic matter in pore waters. McArthur (1985) suggested that variations in lattice-bound CO2 concentrations in fluorapatite are the result of variable pH buffering in pore waters. Romankevich and Baturin (1972) and Price and Calvert (1978) showed that for Recent south west African phosphorites increases in fluorapatite correlate with increases in fluorapatite cementation and in phosphorus and fluorine uptake, and LeGeros et al. (1967) demonstrated that carbonate substitution limited the size and acicularity of synthetic apatite crystallites. McArthur (1978) and McLennan (1980) concluded that carbonate is expelled from the francolite lattice during weathering, metamorphism or diagenesis. Fig. 14 illustrates that there is a general trend towards increasing lattice-bound carbonate contents (as CO2%) with decreasing d13C values in the Duwi phosphorites. This trend can be attributed to a progressive increase in carbonate substitution in the fluorite lattice in response to increase to increased pore-water carbonate ion concentrations as the result of organic matter decomposition. Higher CO2 and F contents in the francolite of the phosphatic mudclasts compared with the authigenic nodules of Peru and Chile margins, and difference in trend in CO2% and d13C binary plot in different localities and horizons indicate that the chemical composition of the phosphatic grains of the Duwi phosphorites has been changed after formation during diagenesis by taking CO2 and F into the structure from the surrounding pore-water. Glenn and Arthur (1990) pointed out that the formation temperature of Egyptian phosphorites may be around

25 C, while those on the Peru shelf are generally less than 14 C based on oxygen isotopic studies of phosphorites. This implies that the low d18O values can be attributed to a decrease in the isotopic composition by meteoric water mixing rather during diagenesis than the increased temperatures conditions. McArthur et al., 1987) interpreted a similar d18O values (4.5& to 9.5&) with a positive correlation with d13C in the Upper Varswater Formation, South Africa, as mainly resulting from alteration of francolite by the action of meteoric waters. Shemesh et al. (1983) and McArthur et al. (1986) interpreted the 18O depletion in ancient phosphorites such as Jordan, South Africa, Phosphoria Formation as a result of post-deposition alterations. The RCO2 in the interstitial water system is inherited from dissolution of marine carbonate with additional contribution from atmospheric CO2 and oxidation of organic materials. d13C values of dissolved carbonate in sea water and marine carbonate fall in relatively narrow range between 0.5& and +0.5& (Kroopnick, 1985) whilst the values d13C values of organic materials are generally more depleted in 13C in the range between 18& and 24& (Anderson and Arthur, 1983). Therefore, the negative and variable d13C values of Duwi phosphorites suggest that relatively large percentage of carbonate ions in these phosphorites might have been derived from oxidation of organic carbon. The linear relationship between d13C and CO2 content for the studied phosphorites (Fig. 14) supports this interpretation as it was previously discussed. Carbon and oxygen isotope data, thus, indicate that the chemical composition of these phosphorites changed during diagenesis via degradation of organic matter and/or exposure to meteoric water. Spatial differences in the d13C and d18O plot between phosphorites from Red Sea and Nile Valley and Abu-Tartur areas confirm this interpretation and explain the effect of diagenetic conditions on the compositional changes where the CO2 reflects the conditions of depositional places which have some differences and does not reflect the formational place which supposed to be same. The presence of SO2 4 in francolite lattice as an isomorphous substitution of PO3 has been suggested by many 4 researchers, and is now generally accepted (Gulbrandsen, 1969). The degree of substitution depends on the SO2 4 content of the aqueous phase that was in equilibrium with the phosphorites during their formation, i.e. high SO2 4 content in the francolite indicate high SO2 content in the forma4 tional environment (McArthur, 1978). The difference in the SO3 contents in the phosphorites from different horizons and localities means that SO3 distribution is controlled by localities and horizons, which in turn with the higher SO3contents in these phosphorites (average = 1.94%) compared to the authigenic phosphatic nodules of Peru margin (0.99%) proposed by Baturin and Bezrukov (1979), support the interpretation that chemical composition of the Duwi phosphorites has been changed after deposition during diagenesis.

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

6. Conclusions Upper Cretaceous phosphorites in Egypt were subjected to detailed mineralogical, geochemical, and isotopic analyses to examine their genesis and diagenesis. The compositional homogeneity of the mudclasts, Fe and S zoning in some of bone fragments, differences in the Chemical Index of Alteration between the acid-insoluble residues of the phosphorites and the associated rocks, and low Mg content in the phosphatic grains suggest that the phosphatic grains in the Duwi Formation are derived from pre-existing authigenic phosphorites formed outside the depositional basin, which reworked and concentrated in the depositional basin afterward. Higher CO2, SO3 and F contents compared to the recent authigenic phosphorites and negative d13C and d18O values of structural CO2 indicate the importance of diagenesis in the modification of chemical composition of phosphatic grains and the studied apatite was francolitized during diagenesis. Acknowledgements Authors thank Dr. Abuelhassan and Mr. E. Yoshida, S. Yamamoto and Y. Matsushita for their kind help either in the field or laboratory. All thanks to the Ministry of Culture and Education of Japan (MONBUSHO) for the financial support to H.M. Baioumy. References Anderson, T.F., Arthur, M.A., 1983. Stable isotopes of oxygen and carbon and their application to sedimentologic and paleoenvironmental problems. In: Arthur, M.A., (Organizer), Stable Isotopes in Sedimentary Geology. Soc. Econ. Paleontol. Mineral., Short Course No. 10, pp. 1–151. Baioumy, H.M., Tada, R., 2005. Origin of Upper Cretaceous phosphorites in Egypt. Cretaceous Res. 26, 261–275. Baturin, G.N., Bezrukov, P.L., 1979. Phosphorites on the sea floor and their origin. Marine Geol. 31, 317–332. Claypool, G.E., Kaplan, I.R., 1974. The origin and distribution of methane in marine sediments. In: Kaplan, I.R. (Ed.), Natural Gases in Marine Sediments. Plenum, New York, p. 240p. Donnelly, T.H., Shergold, J.H., Southgate, P.N., Barnes, C.J., 1990. Events leading to global phosphogenesis around the Proterozoic– Cambrian transition. In: Notholt, A.J.G., Jarvis, I. (Eds.), Phosphorite research and development. Jour. Geol. Soc. London. Spl. Pub., 52, pp. 273–287. Epstein, S., Buchsbaum, R., Lowestam, H.A., Urey, H.C., 1953. Revised carbonate-water temperature scale. Geol. Soc. Am. Bull. 62, 417–426. Glenn, C.R., 1990. Depositional sequence of the Duwi, Sibaiya and Phosphate Formations, Egypt: phosphogenesis and glauconitization in a late Cretaceous epeiric sea. In: Notholt A.J.G., Jarvis, I., (Eds.), Phosphorite Research and Development. Geol. Soc. Spec. Publ. 52, pp. 205–222. Glenn, C.R., Arthur, M.A., 1990. Anatomy and origin of a Cretaceous phosphorite-greensand giant, Egypt. Sedimentology 37, 123–154. Glenn, C.R., Arthur, M.A., Yeh, H.W., Burnett, W.C., 1988. Carbon isotopic composition and lattice-bound carbonate of Peru-Chile Margin phosphorites. Marine Geol. 80, 287–307.

27

Gulbrandsen, R.A., 1969. Physical and chemical factors in the formation of apatite. Econ. Geol. 64, 365–382. Gulbrandsen, R.A., 1970. Relation of carbon dioxide content of apatite of the phosphoria formation to regional facies. U.S. Geol. Surv. Prof. Paper 700B, 9–13. Hudson, D.J., 1977. Stable isotope and limestone lithification. J. Geol. Soc. London. 133, 637–660. Irwin, H., Curtis, C.D., Coleman, M., 1977. Isotopic evidence for source of diagenetic carbonates formed during burial of organic-rich sediments. Nature 269, 209–213. Jarvis, I., Burentt, W.C., Nathan, Y., Almbaydin, S.M.F., Attia, A.K.M., Castro, L.N., Flicoteaux, R., Hilmy, M.E., Husain, V., Qutawah, A.A., Serjani, A., Zanin, Y.N., 1994. Phosphorite geochemistry: State-of-the-art and environmental concern. In: Concepts and Controversies in Phosphogenesis. Proceeding of the Symposium and Workshop held on 6–10 September, 1993 Switzerland, pp. 643–700. Kolodny, Y., Luz, B., 1992. Phosphate deposits, formation and diagenetic history. In: Clauer, N., Chaudhuri, S. (Eds.), Isotopic Signatures and Sedimentary Records. Springer Verlag, pp. 69–122. Kroopnick, P.M., 1985. The distribution of 13C of RCO2 in the world oceans. Deep-Sea Res. 32, 57–84. LeGeros, R.Z., Trautz, O.R., LeGeros, J.P., Edward, K., 1967. Apatite crystallites-effects of carbonate on morphology. Science 155, 1409– 1411. McArthur, J.M., 1978. Systematic variations in the contents of Na, Sr, CO3, and SO4 in marine carbonate-fluorapatite and their relation to weathering. Chem. Geol. 21, 89–112. McArthur, J.M., 1985. francolite geochemistry-compositional controls during formation, diagenesis, metamorphism and weathering. Geochim. Cosmochim. Acta 49, 23–35. McArthur, J.M., Coleman, M.L., Bremner, J.M., 1980. Carbon and ocygen isotopic composition of structural carbonate in sedimentary francolite. J. Geol. Soc. London 137, 669–673. McArthur, J.M., Benmore, R.A., Coleman, M.L., Soldi, C., Yeh, H.W., O’Brien, G.W., 1986. Stable isotopic characterization of francolite formation. Earth Planet. Sci. Lett. 77, 20–34. McArthur, J.M., Hamilton, P.J., Greensmith, J.T., Boyce, A.J., Fallick, A.E., Birch, G., Walsh, J.N., Benmore, R.A., Coleman, M.L., 1987. Phosphorite geochemistry: Isotopic evidence for meteoric alteration of francolite on a local scale. Chem. Geol. 65, 415–425. McClellan, G.H., Lehr, J.R., 1969. Crystal chemical investigation of natural apatite. Am. Mineral. 54, 1374–1391. McLennan, S.M., 1980. Geochemistry of Archean shale from the Pilbara Supergroup, western Australia. Geochim. Cosmochem. Acta 47, 1211– 1222. Murray, R.W., Jpnes, D.L., Buchholtz, T.B., 1992. Diagenetic formation of bedded chert: evidence from chemistry of the chert-shale couplet. Geology 20, 271–274. Nesbitt, H.W., Young, G.M., 1982. Chemical processes affecting alkalies and alkaline earth during continental weathering. Geochim. Cosmochem. Acta. 44, 1659–1666. Notholt, A.J.G., 1985. Phosphorite resourses in the Mediterranean (Tethyan) phosphatic province: a progress report. Sci. Geol. Mem. 77, 9–21. Price, N.B., Calvert, S.E., 1978. The geochemistry of phosphorites from the Namibian Shelf. Chem. Gfeol. 23, 151–170. Romankevich, Y.A., Baturin, G.N., 1972. Composition of the organic matter in phosphorites from the continental shelf of Southwest Africa. Geochem. Int. 9, 464–470. Sandberg, S., 1983. An oscillating trend in phanerozoic non-skeletal carbonate mineralogy. Nature 305, 19–22. Scuffert, J.D., Kastner, M., Emanuele, G., Janhnke, R.A., 1990. Carbonate-ion substitution in francolite: a new equation. Geochim. Cosmochim. Acta 54, 2323–2328. Shemesh, A., Kolodny, Y., Luz, B., 1983. Oxygen isotopes in phosphate of biogenic apatites, II. Phosphorite rocks. Earth Planet. Sci. Lett. 64, 405–416.

28

H.M. Baioumy et al. / Journal of African Earth Sciences 49 (2007) 12–28

Soudry, D., 1992. Primary bedded phosphorites in the Campanian Mishash Formation, Negev, Southern Israel. Sediment. Geol. 80, 77–88. Spanderashvilli, G.I., Mansour, M., 1970. The Egyptian phosphorites. In: Moharram, O. et al., (Eds.), Studies on Some Mineral Deposits of Egypt, pp. 89–106.

Urey, H.C., Lowenstam, H.A., Epstein, S., McKinney, M.H., 1951. Measurements of paleotemperatures and temperatures of the Upper Cretaceous of England, Denmark and the south-eastern U.S. Geol. Soc. Am. Bull. 62, 399–416.