Paleogene succession in north Eastern Desert, Egypt

Journal of African Earth Sciences 81 (2013) 35–59

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Stratigraphy, sedimentology and tectonic evolution of the Upper Cretaceous/Paleogene succession in north Eastern Desert, Egypt Abdalla M. El Ayyat ⇑, Nageh A. Obaidalla Geology Department, Faculty of Science, Assiut University, Assiut 71516, Egypt

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Article history: Received 9 February 2012 Received in revised form 22 December 2012 Accepted 8 January 2013 Available online 8 February 2013 Keywords: North Eastern Desert Upper Cretaceous/Paleogene Syn-depositional tectonic events Syrian Arc System Facies belts Slope to basin transect

a b s t r a c t The stratigraphy, sedimentology and syn-depositional tectonic events (SdTEs) of the Upper Cretaceous/ Paleogene (K–P) succession at four localities in north Eastern Desert (NED) of Egypt have been studied. These localities are distributed from south-southwest to north-northeast at Gebel Millaha, at North Wadi Qena, at Wadi El Dakhal, and at Saint Paul Monastery. Lithostratigraphically, four rock units have been recorded: Sudr Formation (Campanian–Maastrichtian); Dakhla Formation (Danian–Selandian); Tarawan Formation (Selandian–Thanetian) and Esna Formation (Thanetian–Ypresian). These rock units are not completely represented all over the study area because some of them are absent at certain sites and others have variable thicknesses. Biostratigrapgically, 18 planktonic foraminiferal zones have been recorded. These are in stratigraphic order: Globotruncana ventricosa Zone (Campanian); Gansserina gansseri, Contusotruncana contusa, Recimguembelina fructicosa, Pseudohastigerina hariaensis, Pseudohastigerina palpebra and Plummerita hantkenenoides zones (Maastrichtian); Praemurica incostans, Praemurica uncinata, Morozovella angulata and Praemurica carinata/Igorina albeari zones (Danian); Igorina albeari, Globanomanlina pseudomenradii/Parasubbotina variospira, Acarinina subsphaerica, Acarinina soldadoensis/ Globanomanlina pseudomenardii and Morozovella velascoensis zones (Selandian/Thantian); and Acarinina sibaiyaensis, Pseudohastigerina wilcoxensis/Morozovella velascoensis zones (earliest Ypresian). Sedimentologically, four sedimentary facies belts forming southwest gently-dipping slope to basin transect have been detected. They include tidal flats, outer shelf, slumped continental slope and open marine hemipelagic facies. This transect can be subdivided into a stable basin plain plus outer shelf in the extreme southwestern parts; and an unstable slope shelf platform in the northeastern parts. The unstable slope shelf platform is characterized by open marine hemipelagic, fine-grained limestones and fine siliciclastic shales (Sudr, Dakhla, Tarawan and Esna formations). The northeastern parts are marked by little contents of planktonic foraminifera and dolomitized, slumped carbonates, intercalated with basinal facies. Tectonically, four remarkable syn-depositional tectonic events (SdTEs) controlled the evolution of the studied succession. These events took place strongly within the Campanian–Ypresian time interval and were still active till Late Eocene. These events took place at: the Santonian/Campanian (S/C) boundary; the Campanian/Maastrichtian (C/M) boundary; the Cretaceous/Paleogene (K/P) boundary; and the Middle Paleocene–Early Eocene interval. These tectonic events are four pronounced phases in the tectonic history of the Syrian Arc System (SAS), the collision of the Afro-Arabian and Eurasian plates as well as the closure of the Tethys Sea. Ó 2013 Elsevier Ltd. All rights reserved.

1. Introduction The K–P succession covers about one-third of the total surface area of the Egyptian territory (Said, 1990). It is represented by different marine facies (shale, chalk and marl with limestone interbeds). Thick section of limestone and dolomite (the Thebes Formation) covers the whole succession. The study area is located in the NED and delineated by latitude 28°300 and 30°000 N and ⇑ Corresponding author. E-mail addresses: [email protected], [email protected] (A.M. El Ayyat), [email protected] (N.A. Obaidalla). 1464-343X/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.jafrearsci.2013.01.007

longitudes 32°150 and 33°000 E, west of the Gulf of Suez (Fig. 1). It exhibits good exposures of K–P rocks and is considered as a part of the mobile shelf of Egypt (Said, 1990). These rocks crop out as a single great plateau (e.g. the Southern Galala plateau). The dissection of this plateau by several faults has resulted in the exposure of rock succession along most of the plateau. Folds and faults trending northeast are known on the tectonic map of the study area (Mazhar et al., 1979). Previous studies dealing with the area or its environs include the works of Haggag (1991), Keheila (2000), Hussein and Abd-Allah (2001) and El Ayyat and Obaidalla (2005). One of the major problems encountered by all investigators is the poor age control of the

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Fig. 1. Location map showing the study area and the locations of the measured sections.

sequence under study. This has provided little or no information on the timing of stratigraphic events and thus confused regional as well as global biostratigraphic correlation. Table 1 summarizes the lithostratigraphic correlations of the different proposed rock units of K–P exposures under study and the surrounding areas. The aims of this study are to determine the lithological characteristics of a number of key outcrops in the area studied; redefine and correlate the Maastrichtian–Ypresian lithostratigraphic subdivisions in northeastern Egypt from Saint Paul Monastery in the north to Gebel Millaha in the south; laying out high resolution planktonic foraminiferal biostratigraphy for accurate age dating and locating the different hiatuses interrupting the sedimentation regime; provide a comprehensive view of the sedimentological model as well as reconstruct the depositional history of the slope to basin transect for the study area. The objectives also include understanding the superposition of different tectonic events, the style of tectonic deformation related to the impact of SAS on sedimentation in the area under consideration. To achieve the previous goals; four stratigraphic sections, representing most of the rock units, have been measured, described, and sampled in detail. They have been subjected to detailed field examinations to recognize and interpret the lithological changes, trace fossils, bioturbation and erosion surfaces. Samples have been

processed following the standard micropaleontological techniques. The samples have been disaggregated in water and washed through a 63 lm sieve. This procedure was repeated until foraminiferal tests were recovered with clean surface texture. All specimens have been picked, identified and mounted on microslides for permanent record. Thin sections and cut slabs have been prepared from the hard rock varieties. Detailed petrographic analyses such as point counting of the ortho-and allo-chemical components, carbonate depositional textures following Dunham (1962) and Embry and Klovan (1971), and water energy index following the methods described by Leighton and Penedexter (1962) and Plumley et al. (1962) have been carried out (Figs. 2–6). Paleoecologic interpretation and basin evolution have been attempted based on the fossil content, litho- and microfacies associations, tectonic behavior of the region and previous related studies. Information from K–P successions in adjacent areas have been integrated and compared to elucidate the paleogeographic setting of the studied basin and to reconstruct a more precise depositional history. Staining technique of Katz and Friedman (1965) is used to differentiate the dolomites. The available data have been used to correlate the studied sections and to illustrate the large scale depositional architecture of the sedimentary sequences. These data form the basis of our regional interpretation of uplift, erosion, sedimentation and

Table 1 Correlation chart of the rock units advocated in this study with those proposed by different authors for the Upper Cretaceous/Paleogene successions at various localities in north Eastern Desert, Egypt. (See below-mentioned references for further information.)

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Fig. 2. Legend for symbols used in the present study.

renewed subsidence by throwing more light on the SdTE, which affected the studied sequence and formulated the sedimentary basins.

2. Geological setting K–P time witnessed drastic geologic and paleo-biologic events all over the world. This is also the case in northern Egypt, where tectonic movements are thought to have occurred during this interval. Such movements are referred as SAS and they seem to have influenced the sedimentation of K–P deposits in the region (Bartov and Steinitz, 1977). Around the Gulf of Suez, these movements are masked by the more intense and recent tectonism, which led to the generation of the Gulf of Suez rift during Late Eocene and younger times. Major rift-related and rotated fault blocks characterize the present-day geomorphology of plateaus, plains and wadis (Mazhar et al., 1979). The SAS can be traced from Syria to the central Western Desert of Egypt, via Sinai and the NED. It includes the Palmyride and SinaiNegev fold belts, both having similar lithologic and structural characteristics (Shahar, 1994). The Galala plateaus represent a major branch of the SAS in the Eastern Desert, which are characterized by Late Cretaceous uplift in the north and subsidence farther to the south. In contrast to various SAS localities in Egypt and neighboring areas, the Galala plateaus exhibit a unique K–P carbonatesiliciclastics platform to slope succession. This is comparable to other island-like Tethyan elevations rimmed by carbonate platforms, such as the Maiella in Italy (Elberi et al., 1994), and Sicily (Camion et al., 1988). The complex uplifts and domal anticlines of the SAS were formed during the closure of the Neo-Tethys (Stampfli et al., 1995), as a consequence of the convergence of the African and Eurasian plates. Northeastern Egypt, situated at the northern edge of the African–Arabian Craton, was affected during K–P times by east-northeast oriented dextral wrench faulting. This resulted in transgressive movements and the inversion of the Late Triassic half grabens that cut east-northeast across the northern rim of the African–Arabian plate. The area under consideration can be divided into three east-northeast striking facies belts: (1) The northern continental slope deposits that occupy the extreme northern parts of the study area (e.g. Saint Paul). By this position, they can be

considered as a transition zone between the carbonate shelf platform to the north outside the study area, and the central sub-basin (the Southern Galala Sub-basin, Kuss et al., 1999). They are characterized by Upper Cretaceous deposits. (2) The central sub-basin, which was formed in the Late Campanian by the Northern Galala/Wadi Araba Uplift (NG/WAU), it represents the northern part of the Eastern Desert intrashelf basin. Sedimentary wedges, prograding southward from the NG/WAU, led to an increase in loading and subsidence in the sub-basin. The sub-basin configuration became apparent in the Paleocene, as shown by paleobathymetric estimates (Scheibner et al., 2000). (3) The southern intrashelf shallower basin, which occupies the southern portion of the study area and filled with K–P rocks. The interbedding of open marine pelagic facies with shallower outer shelf deposits is a common feature of this intrashelf basin. The east-northeast trending NG/WDU influenced the K–P sedimentation processes in the area under study. Consequently, they are characterized by a gently south-dipping platform that rims the uplift in the north and interfingers with transitional slope sediments and hemipelagic deposits of the sub-basin farther south. Paraconformities and discontinuities of long duration occur at various levels within each section (Kuss et al., 1999). Discontinuities of much shorter duration in slope areas are indicated by the penecontemporaneous reworking of sediments. The same is true in areas that are more basinal where the absence of individual subzones suggests an incomplete stratigraphic record.

3. Lithostratigraphy 3.1. Sudr Formation The Sudr Formation was first introduced by Ghorab (1961) to describe a thick chalk and chalky limestone (247 m thick) at Wadi Sudr in west-central Sinai (type locality). It represents the oldest exposed rock unit (Plate A5) in the study area (Campanian–Maastrichtian). It is made up of rhythmic chalk and chalky limestone (40–80 m thick) interbedded with slightly wavy, very thin beds or lamina of greenish grey calcareous shales (4–10 cm thick). The formation is poor in megafossils, however, the basal part at North Wadi Qena, includes crowds of Pychnodonte, Pecten and Baculites taxa. The formation is rich in microfossils, particularly planktonic

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Fig. 3. Vertical distribution of the main sedimentological characteristics, microfacies and their depositional sites for the different rock units at Saint Paul locality (legend in Fig. 2).

foraminifera (Figs. 3–6). Slump and slide sedimentary structures are common with various scales. These structures show increase in abundance and magnitude within the lower part at both Saint Paul and Wadi El Dakhal localities. The Sudr Formation exhibits a considerable variation in thickness at different localities (Fig. 7 and Table 2). The lower contact with the Matulla Formation is

unconformable (Plate A1). The upper contact of the Sudr Formation with the overlying Dakhla Formation is a regional and widespread unconformity (K–P boundary). Soliman et al. (1986) believed that the Dakhla Formation changes laterally in north Egypt into chalk unit (Sudr Chalk) in NED and Khoman Chalk in northwestern Desert.

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Fig. 4. Vertical distribution of the main sedimentological characteristics, microfacies and their depositional sites for the different rock units at Wadi El Dakhal locality. (legend in Fig. 2). Tar. Fm. = Tarawan Formation; E. F. = Esna Formation; T. F. = Thebes Formation.

3.2. Dakhla Formation The term Dakhla Shale was first introduced by Said (1962) to describe the unit overlying the Duwi Phosphate and underlying the Tarawan Chalk. Awad and Ghobrial (1965) used the term Dakhla Formation in the Kharga area. In the study area, the Dakhla Formation exhibits a considerable variation in thickness (Fig. 7 and

Table 2). It consists of a thick succession of light olive grey and dark greenish grey shales (Plate A4) with intercalated argillaceous, phosphatic limestone bands. The shales are enriched with planktonic and benthonic foraminifera (Figs. 3–6). The Dakhla Formation overlies unconformably the Sudr Formation and underlies unconformably the Tarawan Formation (Plate A2). At extreme northnorthwest (Saint Paul locality), the Dakhla Formation underlies

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Fig. 5. Vertical distribution of the main sedimentological characteristics, microfacies and their depositional sites for the different rock units at Gebel Millaha locality (legend in Fig. 2).

unconformably the Thebes Formation (Plate A3) with complete absence of both Tarawan and Esna formations. 3.3. Tarawan Formation The Tarawan Chalk has been introduced by Awad and Ghobrial (1965) to designate a chalky limestone, containing bands of

marl. Here, the Tarawan Formation is highly jointed and fractured, forming steep standing face at the top of the plateau. High abundance of planktonic forams and absence of coarse detrital influx in the sediment indicate an open marine condition of moderate depth. This rock unit displays gradual changes in thickness at the extreme north-northeast direction (Saint Paul locality), where it missed completely (Fig. 7 and Table 2).

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Fig. 6. Vertical distribution of the main sedimentological characteristics, microfacies and their depositional sites for the different rock units at North Wadi Qena locality (legend in Fig. 2).

The base of the Tarawan Formation marks a hiatus with the underlying Dakhla Formation, due to major regression of the sea during Middle Paleocene. This basal part is piped by Callianassid burrows and riddled with water-worn bioclasts. As a consequence, a bioturbated conglomeratic bed with phosphatic pebbles, vertebrate remains and reworked fauna, at the base,

marks a hiatus between Dakhla and Tarawan formations. On the other hand, conformable relationships are recorded between the Tarawan and the overlying Esna Formation, where the marly and argillaceous chalky limestone of the Tarawan passes gradually into the calcareous shales of the Esna Formation (Fig. 7).

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Fig. 7. Lithostratigraphic correlation of the studied rock units from NNE to SSW direction.

Table 2 Vertical thickness (in meters) of the studied rock units at different localities. Rock unit

G. Millaha

N. Wadi Qena

W. El Dakhal

Saint Paul

Esna formation Tarawan formation Dakhla formation Sudr formation

14.5 3 6.5 23

27.5 2.5 18.5 40.5

3 2 4.5 103

0 0 4.5 69.5

3.4. Esna Formation The term Esna Shale has been used firstly by Beadnell (1905) to define the laminated green and grey shale exposed at Gebel Oweina, opposite Esna, Nile Valley. In the study area, it overlies the Tarawan Formation and underlies the Eocene carbonates of the Thebes Formation (Plate A5). It occurs as dissected and badly exposed, slope forming calcareous shales, being almost obscured by the scree of the overlying Eocene rocks. The depositional basin of the Esna Formation has the same paleogeography of the underlying Tarawan Formation (Fig. 7). It dies toward north-northeast where it disappears completely at Saint Paul. It is flooded with planktonic and benthonic foraminifera and moderately bioturbated at certain levels. Travelling south (to North Wadi Qena and Gebel Millaha), the upper part is slightly dolomitized (Figs. 5 and 6). Contacts with the overlying Thebes Formation are conformable at North Wadi Qena and unconformable at both Wadi El Dakhal and Gebel Millaha localities (Fig. 7).

planktonic foraminiferal zones are arranged from base to top as follows:

4.1. Globotruncana ventricosa Zone This zone has been defined by Dalbiez (1955) as an interval zone from the Lowest Occurrence (LO) of the nominate taxon to the LO of Globotruncanita calcarata (Cushman), to mark the middle part of the Campanian age. It is equivalent to that of Caron (1985). The G. ventricosa Zone is represented all over the studied sections (Figs. 8–10). It is unconformably overlain by the Gansserina gansseri Zone of the Early Maastrichtian age. This reflects a hiatus between the Campanian and Maastrichtian ages (Table 3). This hiatus led to missing of the upper parts of the Campanian and the lower parts of the Maastrichtian successions.

4.2. Gansserina gansseri Zone Brönnimann (1952) defined the G. gansseri Zone as a partial range of the nominate taxon between its LO and the LO of the Abathomphalus mayaroensis (Bolli). Li and Keller (1998) defined this zone to cover the interval from the LO of the G. gansseri (Bolli) to the LO of Contusotruncana contusa (Cushman). Here, the A. mayeroaensis taxon is absent, so that the definition of Li and Keller (1998) to the upper boundary of G. gansseri is here applied. It is overlain conformably by the C. contusa Zone of Middle Maastrichtian age and is well represented all over the studied area.

4. Planktonic foraminiferal zonation The measured sections yield rich and well diversified planktonic foraminiferal fauna. Depending on the stratigraphic range of the index planktonic foraminiferal species, 18 zones are defined (Figs. 8–11) ranging in age from Late Cretaceous (Campanian– Maastrichtian) to Early Paleogene (Danian–Ypresian). The Late Cretaceous planktonic foraminiferal zones of Caron (1985) and Li and Keller (1998) and the Early Paleogene zones of Berggren and Pearson (2005) are here applicable. Table 3 illustrates the planktonic foraminiferal zonation and datum levels of the present study correlated with well-dated international zonations. The proposed

4.3. Contusotruncana contusa Zone Li and Keller (1998) defined the C. contusa Zone as a partial range zone from the LO of the nominate taxon and the Highest Occurrence (HO) of the Globotruncana linneniana (Bolli). In the present study, the G. linneniana (Bolli) species is rarely represented. Accordingly, the upper boundary of this zone is defined at the LO of Recemiguembelina fructicosa taxon and is well represented in the studied sequence (Table 3). The C. contusa Zone is overlain conformably by the R. fructicosa Zone of Late Maastrichtian age.

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Plate A. (1): Field view showing the lower contact of the Sudr Formation with the Matulla Formation. Note the irregular relationship and the strong deformation of the uppermost part of the Matulla Formation. Saint Paul locality. (2): Field view showing the contact between Dakhla Formation and the overlying Tarawan Formation. Strong bioturbation filled with coarser materials. Gebel Millaha locality. (3): Field view showing the unconformity between Dakhla and Thebes Formation. White caliche horizon separating between the two rock units. Saint Paul locality. (4): Field photo showing a thick succession of dark greenish grey shales of the Dakhla Formation. Wadi El Dakhal locality. (5): Panorama view showing complete stratigraphic succession for the studied rock units. North Wadi Qena locality. (6): Polished cut slab showing solution seams developed at bedding planes in the limestone of the Sudr Formation. Wadi El Dakhal locality. (7): Field view showing the Dababiya Quarry Member at the P/E boundary at north Wadi Qena section. (8): Field view showing the Qreiya Beds at the D/S boundary at north Wadi Qena section.

4.4. Recemiguembelina fructicosa Zone The R. fructicosa Zone is also defined by Li and Keller (1998) to define the interval from the LO of the nominate taxon to the LO of the Pseudoguembelina hariaensis Nedergragt. This zone is defined all over the area and is overlain conformably by the P. hariaensis Zone of Late Maastrichtian age. 4.5. Pseudoguembelina hariaensis Zone The absence of A. mayeroaensis taxon (Late Maastrictian marker) in several geographic provinces in the world of shallow continental

shelf setting, led some authors to use another taxa to make highresolution biostratigraphy for the Upper Maastrichtian sediments (Keller, 1988; Pardo et al., 1996; Luciani, 1997). Nedergate (1990) recorded the P. hariaensis taxon to mark the interval of 15 m below the K/P boundary at El Kef, Tunisia. Li and Keller (1998) used this taxon to cover the Upper Maastrichtian sediments from the LO of this zone to the HO of G. gansseri (Bolli) at northwestern Tunisia. Here, this zone is recorded all over the study area and is overlain unconformably by Praemurica uncinata Zone of the Late Danian age at Gebel Millaha section. On contrary, it is overlain conformably by the P. palpebra Zone of latest Maastrichtian age at North Wadi Qena, Wadi El Dakhal and Saint Paul sections.

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Fig. 8. Range chart of the index planktonic foraminifera species, Saint Paul section.

4.6. Pseudoguembelina palpebra Zone

4.7. Plummerita hantkeninoides Zone

It has been defined by Li and Keller (1998) to mark the interval from the HO of G. gansseri (Bolli) to the LO of Plummerita hantkeninoides (Brönniman). Here, this zone is absent at Gebel Millaha, but it is recorded at other localities. It is overlain conformably by the P. hantkeninoides Zone of latest Maastrichtian age.

Pardo et al. (1996) defined the P. hantkeninoides Zone as the total range of the nominate taxon to mark the uppermost part of the Maastrichtian deposits at Agost, Spain. It was also used by Molina et al. (1998), Luciani (2002) and Obaidalla (2005). Here, this zone is absent at Gebel Millaha and occurs in the remainder sections. It is

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Fig. 9. Range chart of the index planktonic foraminifera species, Wadi El Dakhal section.

overlain unconformably by the Praemurica inconstans Zone of earliest Danian age. 4.8. Praemurica inconstans (P1c) Zone The P. inconstans (P1c) Zone was defined by Berggren et al. (1995) as the partial range of the nominate taxon from its LO to the LO of P. uncinata (Bolli). It represents the first Danian plank-

tonic foraminiferal zone at North Wadi Qena, Wadi El Dakhal and Saint Paul sections. On the other hand, this zone is absent at Gebel Millaha section. This indicates a long-lasting hiatus between Cretaceous and Paleogene at Gebel Millaha, which led to the absence of latest Maastrichtian P. palpebra and P. hantkeninoides zones and the earliest Danian Guembelitria cretacea (P0), P. eugubina (Pa), P. pseudobulloides (P1a), S. triloculinoides (P1b) and P. inconstans (P1c) zones (Table 3). Moreover, this hiatus is characterized by

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Fig. 10. Range chart of the index planktonic foraminifera species, Gebel Millaha section.

short duration and led to the absence of the earliest Danian zones at North Wadi Qena, Wadi El Dakhal and Saint Paul. Generally, the hiatus is very well-known in Egypt and shows remarkable variations from one locality to another (e.g. Abdel-Kireem and Samir, 1995; Obaidalla and Kassab, 2000; Obaidalla et al., 2006). The P. inconstans Zone is overlain conformably by the Praemurica uncinata Zone of Late Danian age. 4.9. Praemurica uncinata (P2) Zone This zone was defined by Bolli (1966) to define the interval from the LO of the nominate taxon to the LO of Morozovella angulata (White). This definition was used by Berggren et al. (1995) and Berggren and Pearson (2005). The P. uncinata Zone is recorded at North Wadi Qena, which is conformably overlain by the M. angulata (P3a) Zone of Late Danian age, whereas at Wadi El Dakhal and Saint Paul, this zone is overlain unconformably by Acarinina subs-

phaerica Zone of Thanetian age. In contrast, this zone is missing at Gebel Millaha (Table 3). 4.10. Morozovella angulata (P3a) Zone Hillebrandt (1966) defined the M. angulata Zone to cover the interval from the LO of the nominate taxon to the LO of I. pusilla (Bolli). Berggren et al. (1995) and Berggren and Pearson (2005) redefined the upper boundary of this zone by the LO of I. albeari (Cushman and Bermüdez) according to the stratigraphic priority of I. Pusilla (Bolli) and I. albeari (Cushman and Bermüdez) respectively. The M. angulata Zone is recorded at North Wadi Qena and Gebel Millaha. It is overlain unconformably by the I. albeari/P. carinata (P3b, lower most part) Zone of latest Danian age and I. albeari (P3b, main part) of earliest Selandian age at North Wadi Qena, which denotes continuous sedimentation between the Danian/ Selandian (D/S) boundary. Concurrently, this zone is overlain

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Fig. 11. Range chart of the index planktonic foraminifera species, north Wadi Qena section.

unconformably by I. albeari (P3b, main part) Zone at Gebel Millaha. This indicates that the D/S boundary is marked by very short hiatus, which led to the absence of I. albeari/P. carinata (P3b, lower most part) of latest Danian age. On the other hand, the D/S boundary is characterized by a long-range hiatus at Wadi El Dakhal and Saint Paul, which led to the absence of latest Danian M. angulata and I. albeari/P. carinata zones and Selandian I. albeari and Globanomalina/P. variospira zones (Table 3).

(P3b, main part) of Early Selandian age (Table 3). Moreover, the D/S boundary at North Wadi Qena is characterized by the occurrence of distinctive lithologic beds (organic-rich phosphatic shale beds) named as Qreyia Beds (Plate A8) by Soliman and Obaidalla (2010). This proves the continuous sedimentation at the Danian/ Selandian (D/S) boundary at this site.

4.11. Igorina albeari/Praemurica carinata (P3a, lower most part) Zone

The I. albeari Zone (P3b) Zone was defined by Berggren et al. (1995) as the partial range of the nominate taxon to cover the interval from the LO of the nominate taxon to the LO of Globanomalina pseudomenardii (Bolli). In the present study, the lower boundary of this zone is defined, following the definition of Obaidalla et al. (2009), at the HO of P. carinata (El Naggar). The I. albeari (P3b, main part) Zone is recorded only at Gebel Millaha and North

This zone was proposed by Obaidalla et al. (2009) as a concurrent range zone to cover the interval of the latest Danian from the LO of the I. albeari (Cushman and Bermüdez) to the HO of P. carinata (El Naggar). The I. albeari/P. carinata Zone is only recorded at North Wadi Qena, and is conformably overlain by I. albeari Zone

4.12. Igorina albeari (P3a, main part) Zone

Table 3 Comparison between the planktonic foraminiferal biozones of the present work with international global schemes (age assignment based on Bralower et al. (1995) for Late Cretaceous and on Berggren and Pearson (2005) for Paleogene.

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Wadi Qena sections and is overlain conformably by Globanomalina pseudomenardii/Parasubbotina variospira (P4a) Zone of Selandian age. 4.13. Globanomalina pseudomenardii/Parasubbotina variospira (P4a) Zone Berggren and Pearson (2005) redefined the subdivisions of the P4 Zone. They (op. cit.) modified the upper boundary of P4a by the HO of P. variospirs (Belford) instead of the LO of Acarinina subsphaerica (Subbotina). This zone is equivalent to the lower part of Berggren et al. (1995) Zone. The Globanomalina pseudomenardii/ Parasubbotina variospira (P4a) Zone is also recorded only at Gebel Millaha and North Wadi Qena sections and is overlain conformably by Acarinina subsphaerica (P4b) Zone of Selandian–Thanetian age. 4.14. Acarinina subsphaerica (P4b) Zone It was defined by Bolli (1966) as the total range of the nominate taxon. Berggren et al. (1995) also defined the P4b Zone as the total range of A. subsphaerica (Subbotina). Recently, Berggren and Pearson (2005) modified the lower and upper boundaries of P4b according to the HO and LO of P. variospira and A. soldadoensis, respectively. The Acarinina subsphaerica (P4b) Zone is defined at Gebel Millaha and North Wadi Qena and is overlain conformably by Acarinina soldadoensis/Globanomalina pseudomenardii (P4c) Zone of Thanetian age. It is also defined at Wadi El Dakhal section and is overlain unconformably by Morozovella velascoensis (P5) Zone of latest Thanetian age (Table 3). 4.15. Acarinina soldadoensis/Globanomalina pseudomenardii (P4c) Zone This zone was defined by Berggren et al. (1995) as a concurrent range zone to cover the interval from the LO of A. soldadoensis (Brönnimann) to the HO of G. pseudomenardii (Bolli). Berggren and Pearson (2005) follow the same bioevents. The A. soldadoensis/G. pseudomenardii (P4c) Zone is recorded only at Gebel Millaha and North Wadi Qena sections and is overlain conformably by M. velascoensis (P5) Zone of latest Thanetian age (Table 3).

earliest Ypressian, which indicates the occurrence of short-range hiatus at the P/E boundary. This hiatus led to the missing of Dababiya Quarry Member that is represented by the A. sibaiyaensis (E1) Zone. On contrary, at Saint Paul and Wadi El Dakhal, the sediments of the latest Thanetian and/or earliest Ypresian are missing due to a long-range hiatus at the P/E boundary. This hiatus led to the absence of M. velascoensis and/or A. sibaiyaensis and Pseudohastigerina wilcoxensis/M. velascoensis zones respectively (Table 3). 4.17. Acarinina sibaiyaensis (E1) Zone Pardo et al. (1999) proposed this zone to define the P/E boundary to cover the interval from the LO of the nominate taxon to the LO of Pseudohastigerina wilcoxensis (Cushman and Ponton). Berggren and Pearson (2005) used this bioevents to define the E1 Zone at the P/E boundary. In this study, the A. sibaiyaensis (E1) Zone occurs only at North Wadi Qena section and is overlain conformably by Pseudohastigerina wilcoxensis/M. velascoensis (E2) Zone of earliest Ypressian age. 4.18. Pseudohastigerina wilcoxensis/Morozovella velascoensis (E2) Zone Molina et al. (1999) defined the P. wilcoxensis Zone to cover the interval from the LO of the nominate taxon to the HO of M. velascoensis (Cushman). Berggren and Ouda (2003a,b) and Berggren and Pearson (2005) used these bioevents to define the P5c and E2 zones, respectively (Table 3). The P. wilcoxensis/M. velascoensis (E2) Zone occurs at Gebel Millaha and North Wadi Qena sections, but it is absent at Wadi El Dakhal and Saint Paul due to a hiatus at the P/E boundary. This zone is considered as the youngest one and terminates the zonation scheme of the area under consideration. 5. Facies analyses and paleoenvironments 5.1. Tidal flat (intertidal to shallow subtidal) facies belt

4.16. Morozovella velascoensis (P5) Zone

5.1.1. Occurrence It is concentrated around Gebel Millaha with a general increase in thickness to south-southwest direction (Fig. 12).

The M. velascoensis (P5) Zone was defined by Bolli (1966) to mark the interval from the HO of G. pseudomenardii (Bolli) to HO of the nominate taxon. Berggren and Ouda (2003a,b) subdivided P5 Zone into three subzones namely: P5a (latest Thanetian), P5b to recognize the Carbon Isotope Excursion interval and P5c to recognize the post CIE interval (earliest Ypressian). Berggren and Pearson (2005) used alphanumeric notation ‘‘P’’ for the planktonic foraminiferal Paleocene zones and alphanumeric notation ‘‘E’’ for the planktonic foraminiferal Eocene zones (Table 3). They (op. cit.) used the P5 Zone for the latest Thanetian and E1 and E2 zones for the earliest Ypressian. According to them, the P5 Zone was defined to cover the interval from the HO of G. pseudomenardii (Bolli) to the LO of A. sibaiyaensis (El Naggar). In Egypt, the Paleocene/Eocene boundary is located at the base of organic-rich phosphatic shale sediments named as Dababiya Quarry Member (Aubry et al., 2007). This member is well represented all over the geographic provinces in Egypt (e.g. Knox et al., 2003; Obaidalla, 2006). The Dababiya Quarry Member (Plate A7) is well represented at North Wadi Qena section, and the P/E boundary lies at the base of this member. At this section, the M. velascoensis (P5) is conformably overlain by A. sibaiyaensis (E1) Zone of earliest Ypressian age (Table 3). At Gebel Millaha, this zone is unconformably overlain by Pseudohastigrina wilcoxensis/M. velascoensis (E2) Zone of the

5.1.2. Description Prevailing rock types are mainly marly phosphatic limestones (shallow subtidal) and/or dolomite and dolomitic limestone (intertidal). Occasionally, differentiation between intertidal and shallow subtidal is unclear due to much lithological similarities. Vertical and horizontal burrows become dominant upwards. Most of bed contacts, however, are characterized by erosional lower bounding surfaces. In each situation of intertidal flat, restriction in fauna is a remarkable feature. Some scattered and fragmented foraminiferal individuals are sparse. In thin section, the rock is composed of dolomite rhombs, which can be differentiated into two types. The first type (70–80% of the rock) is represented by fine-grained rhombs ranging from 10 to 50 lm. The second type (15–20%) is less abundant (Plate B1). Concerning subtidal sediments, it is interstratified with basinal sediments or open marine outer shelf (Fig. 12). It crops out as thin to medium bedded, greyish white to yellowish, chalky limestone, moderately indurated. Horizontal and/or oblique bioturbation is common (Plate B2). It consists of tightly packed skeletal particles, reworked phosphate lithoclasts and vertebrate remains, all embedded in a lime mud matrix. The skeletal particles account for 50– 60% of the total rock and include foraminiferal tests with some thin

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Fig. 12. Lateral and vertical distribution of the sedimentary facies from NNE to SSW for the Upper Cretaceous/Paleogene succession under study in north Eastern Desert.

bivalve debris and ostracods. Most of the forams are of planktonic type with rare benthonics. Generally, the foraminiferal content is low to moderate. The Planktonic/benthonic (P/B) ratio is relatively low (30:70). 5.1.3. Interpretation The petrographic evidences confirm the replacement (secondary) origin for the dolomite in the intertidal flat as evidenced by the different sizes of the dolomite crystals (Scholle and Scholle, 2003). It is believed that during compaction, the swelled smectite which is a major clay mineral component of the interbedded shales, converts into the more stable unswelled illite with release of Mg2+ (McHargue and Price, 1982). The latter is used to form dolomite. Restriction in fauna may reflect stressful environmental condition. Phosphatic packstone beds with erosional bases may reflect high-energy events (i.e. storms). The low to moderate foraminiferal content and low value of P/B ratio indicate that the depositional environment was shallow subtidal suite. The sediments of this facies were accumulated in calm water conditions (EI = I) with short intervals of intermittently agitated water regimes (EI = II, Figs. 3–6). 5.2. Open marine, outer shelf facies belt 5.2.1. Occurrence It is recorded mainly in the central and south-western parts, but it disappears northeast, where it dies completely at Saint Paul locality (Fig. 12). 5.2.2. Description It builds the main bulk lithology of the Tarawan Formation and a major part of the Sudr Formation. It occupies sheet-like bodies of nearly constant thickness. As a principal facies of the Tarawan Formation, it crops out as vertical face of white limestone contrasting with the shales of the underlying and overlying Dakhla and Esna formations, respectively (Plate A5). At certain levels, it is less pure and includes some bands of calcareous shale and marls. Sparse

chert nodules and lenses (5–10 cm diameter) are scattered throughout bedding or arranged in discontinuous bands. The basal part is marked by phosphatic bioturbated limestone, which rests on an irregular erosive surface. The main components are well preserved planktonic foraminifera (20–40%); benthonic foraminifera (5–8%), ostracodes (up to 3%) and fine bioclastics of silt size (up to 15%). The matrix consists of pure and dolomitized micrite (Plate B3). Foraminiferal wackestone and packstone horizons with irregular bases are common. This facies is characterized by a relatively high P/B ratio (73:27). 5.2.3. Interpretation The scarcity of benthonic fossils suggests that this facies was formed as a hemipelagic deposit in deep calm water, in an open marine environment, normally below the range of current activity. The presence of wackestone and packstone beds with erosional lower contacts is inferred to reflect high-energy events that disrupted substrates and deposited planktonic coquina as lags. Features indicative of subaerial exposure are absent, suggesting that these deposits were continually submerged. This facies was deposited under quiet to intermittently water energies (EI = I to II, Plumley et al., 1962; Flügel, 2004). The high foraminiferal content and relatively high P/B ratios indicate that the deposition site was an outer shelf (Murray, 1991). 5.3. Continental slope (upper and lower) facies belt 5.3.1. Occurrence This facies occurs at both Saint Paul and Wadi El Dakhal as branched sheet-like geometries (Fig. 12). Their main bodies are located northward outside the study area. It is interbedded with deep open marine facies toward southwest. 5.3.2. Description According to lithology, sedimentary structures, P/B ratio, and geometry; this facies might be accumulated on continental slope. It could be differentiated into two sub-facies: the upper and lower

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Plate B. (1): Photomicrograph showing two types of dolomite rhombs. Most of which show idiotopic to hypidiotopic fabric with inequigranular texture. The matrix is formed of lime mud. Intertidal facies, Gebel Millaha locality. PPL. (2): Bioturbated foraminiferal wacke to packstone microfacies. Most of foraminiferal tests are fragmented. Shallow subtidal facies, Gebel Millaha locality. PPL. (3): Laminated, partially dolomitized packstone microfacies of open marine outer shelf facies. North Wadi Qena locality. PPL. (4): Field photo showing synsedimentary slump-folded beds. The bedding is intensively deformed, suggesting plastic deformation during transport. The lower part of the Sudr Formation. Saint Paul locality. (5): Photomicrograph showing foraminiferal packstone of the lower continental slope. Sudr Formation. Saint Paul locality. PPL. (6): Photomicrograph showing foraminiferal packstone microfacies of open marine basinal facies. North Wadi Qena locality. PPL.

continental slope (UCS and LCS). The UCS is represented by foraminiferal limemudstone/wackestone and wackestone. Some packstone lenses are recorded at certain levels. The rock is made up of both planktonics and benthonics, randomly embedded in a dense lime mud. The sediments are slightly to moderately bioturbated. The faunal content forms 15–25% of the rock. The rock is marked by a very high P/B ratio (82:18). The components are represented mainly by planktonic foraminifera (40–50%), lithoclasts (20–40%); collophane grains (17–32%); and fragmented bioclasts (14–25%). The matrix includes a mixture of sparite and micrite. The LCS is represented mainly by foraminiferal wackestone and wacke-/packstones (Plate B5). Foraminiferal packstone sheets of very limited areal extension are interbedded with other textures. This sub-facies holds nearly the same faunal association like that of UCS. On contrary, P/B ratio is higher and the faunal diversity is lower. So, the differentiation of this sub-facies is based mainly on

the presence of slump and slide sedimentary structures, which range in thickness between 3.5 and 10.5 m. These structures can be differentiated into syn-sedimentary slump-folded beds (Plate B4); slide blocks rest on large intraformational truncation surface (slump scars) in argillaceous and cherty limemdstone to wackestone and debris flows. The sediments indicate that the direction of slumping is oriented towards the center of basin to south and south-west. 5.3.3. Interpretation The UCS depositional site is suggested for the following criteria: presence of abundant planktonics; relatively small size of the planktonic and benthonics; high P/B ratio; relative dominance of some index species of benthonics, restricted to the upper continental slope as Bathysiphon, Chilostomella (Murray, 1991) and the absence of real shelfal and nearshore fossil communities of shallow

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water habitants and presence of channel lag deposits. These criteria suggest deposition in upper bathyal zone. It is proposed, here, that the syn-sedimentary slump and slide beds developed by down slope, short periods of catastrophic gravitational sliding and slumping on a steep ramp which was subjected to tectonic subsidence and progressive tilting. Extensive syn-sedimentary distortion of bedding was developed by creep in thin to very thin bedded, upper basinal slope peri-platform lime mudstone (Mc Ilreath and James, 1979). Turbidity currents and debris flows appear to be the dominant transport mechanisms for the down slope movements of coarse detritus on carbonate slopes. Quiet to intermittently (EI = I to II) agitated water conditions controlled the deposition of this facies. 5.4. Deep open marine pelagic (basinal) facies belt 5.4.1. Occurrence It represents nearly 70–85% of the rock record with maximum thickness at North Wadi Qena (Fig. 12). 5.4.2. Description It embraces a wide variety of rock types including limestones and fine siliciclastic shales (e.g., Sudr, Dakhla and Esna formations). The shales of the Dakhla Formation are dark greenish grey, flooded with planktonic and benthonic foraminifera. The Sudr Formation crops out as snow white, chalk and chalky limestones (Plate B4), which are mainly limemudstone to wackestone with frequent packstone horizons (Plate B6). Components are planktonic foraminifera (5–20%), benthonics (2–6%) and ostracodes (2–3%). Esna Formation is made up of greyish green calcareous shales enriched with planktonic and benthonic foraminifera. It is moderately bioturbated at certain levels and/or slightly laminated. The upper part is slightly dolomitized southwards (North Wadi Qena and Gebel Millaha). Bioturbated and argillaceous limestone bands intercalate the shales. Generally, these shales become calcareous upsection. 5.4.3. Interpretation Gradual absence of this facies northwards is attributed to the general shallowing in that direction. The sparsely fossiliferous lime mud-, wacke- and packstones reflect deposition in deep open marine setting far from the reach of terrigenous input. Thin- to laminated bedding implies low energy depositional environment. Planktonic foraminifera theoretically could live anywhere in the sea but are most numerous away from shore over deep water due to less competition from benthonic organisms for nutrients in the deeper mass of water. This facies type attests to deposition in an open marine environment below effective wave base and is largely aphotic. This interpretation is adopted from the predominance of planktonic forams (80:20 P/B ratio); presence of subsolution seams; complete absence of shallow water fauna and sedimentary structures). As a whole, the rocks are generally deposited in low energy (EI = I) water conditions with spasmodic, short term intervals of intermittently agitated (EI = II) regimes. 6. The depositional model and historic evolution The K–P rocks were developed on a passive margin basin during the intermediate closing phase of the Neotethys. These rocks were accumulated on a gently south-southwest dipping slope to basin transect, in contrast to a north-dipping Upper Cretaceous platform, evolved in the NED between Gebel Ataqa and Wadi El Dakhal (Scheibner et al., 2000). The basin development was controlled by K–P uplift, which is related to the northwestern differential movements of the African–Arabian plate forming processes of the Red Sea/Gulf of Suez. Youssef et al. (2000) suggested a gentle

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south-dipping Paleocene–Eocene ramp in the Southern Galala plateau and they believed that this ramp was developed as a consequence of the Late Cretaceous SAS. The K–P rocks were accumulated under calm water, open marine environment. The deposition was associated with a marked rise in sea level at the beginning of the Maastrichtian which continued throughout the Ypresian with short periods of regression and non-deposition. The depositional regime is basically characterized by:  Sedimentation took place at moderate to relatively deep sea waters.  Deposition below normal wave base and usually below the water mixing zone.  Dominant planktonic foraminifera with some benthonics.  Sedimentation of extremely fine-grained homogenous material from suspension.  Absence of strictly shallow water fauna.  Presence of subsolution seams and rock fragments (Plate A6) suggesting greater water depths. The following is a detailed description of the depositional conditions prevailed during the accumulation of the K–P sequence and its sedimentary evolution. 1. The sedimentary environments were shallow, partly deeper, marine shelf during the Early Cretaceous up to the end of Turonian. These environments prevailed over the relatively flat and wide continental shelf of the Arabo-Nubian massif. In Senonian time, sedimentation became considerably deeper and pelagic in facies as a result of a major transgression of the Tethys into the shelf. This transgression inundated Wadi Araba uplift (in the north) with shallow sea water. This uplift was an emerged paleohigh during Middle Cretaceous. This paleohigh (Fig. 15), with relative sea level changes, controlled the distribution of the deposited sediments, their thickness and facies types. 2. The open marine incursion was linked with the initiation of a deep subsiding elongated trough within the continental shelf during the Maastrichtian. As a result of basin floor subsidence and relative sea level rise, thick monotonous pelagic carbonate sedimentation (Sudr Formation) of an open and deep marine environments prevailed during the Maastrichtian time. The carbonate sediments were deposited uniformly all over the entire basin. The basin may have been several meters in depth in its depocenter (Fig. 7), which is located in the northern part of the study area around Wadi El Dakhal locality. 3. During Maastrichtian, the northern part of the Arabo/Nubian shelf continued to subside accompanied with a major marine flooding. The marine transgression resulted in the extension of Tethyan waters over vast areas in NED. The widespread marine transgression seems to coincide with eustatic rise in sea level in North Africa (Haq et al., 1987). Here, deposition took place under open marine environment with normal salinity and warm tropical water enriched in planktonic foraminifera. Moreover, the open marine basinal sedimentation still dominant at the basin center. Concurrently, on the paleohighs at the basin margins (e. g. Saint Paul), the basinal sedimentation was intermittent and interbedded upsection with relatively shallower sediments of continental slope or outer shelf origin. 4. At the end of Maastrichtian, a short regression took place. The resulted hiatuses in sedimentation (K/P boundary) have different magnitudes and duration in areas geographically nearby each other. 5. During Danian–Selandian, a major transgression took place, which is attributed to continuous subsidence of the AraboNubian shelf combined with a prominent rise in sea level. Conditions of open marine prevailed once again. These conditions

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favored deposition of a thick monotonous succession of greyish and green shales (Dakhla Formation), locally being calcareous, with varying amounts of planktonic and benthonic forams. It

seems that, the environment of deposition implies a continuation of the same conditions that prevailed during the Maastrichtian. Close inspection of the paleogeography of the Dakhla

Fig. 13. Sketch diagram showing two dimensional (A–D) and block representations (1–4) depicting the evolution of the sedimentary basins for the different rock units under study covering the time interval from Campanian to Early Ypresian.

Fig. 14. Depositional facies model for the Upper Cretaceous/Paleogene sedimentary sequence, north Eastern Desert, demonstrating the distribution of the main ortho-and allochemical components, depositional textures and sedimentary structures.

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Formation basin indicates variation in thickness as well as in age (Fig. 13). The depocenter of sedimentation is traced around North Wadi Qena and pinches out laterally. Moreover, the age of the Dakhla Formation is being younger southward. These variations in thickness and age highlight the effect of SdTE and relative sea level changes on the sedimentation. The great thickness of the Maastrichtian–Danian rocks is an indication of continuous sedimentation in a subsiding basin. Minor breaks in sedimentation mark short periods of submarine hiatus. 6. With the passage of time, the rate of subsidence was relatively decreased and the basin became filled with sediments during Maastrichtian–Danian. The basin fillings led to accumulation of Sudr and Dakhla formations (Fig. 13). When the rate of subsidence is slow and the rate of sediment accumulation is great, the depositional basin was covered with shallow, bioturbated and regressive sediments with relatively low foraminiferal abundance and diversity, deposited over most of the basin, on the top of the Dakhla Formation. Consequently, the former trough of the basin became part of the Arabian continental shelf. Afterwards, the basin started to lose its morphologic expression as an elongated deep subsiding trough during Selandian–Thanetian. The facies architecture reflects the evolution from an old distally-steepened ramp (Maastrichtian–Danian) to a new homoclinal ramp during Selandian–Thanetian. Upon this new ramp, sedimentation of the Tarawan Formation began in an open outer shelf domain. The deposition took place under normal salinity and warm tropical water enriched in planktonic with some benthonic foraminifera. Depositional settings became shallow upwards and then followed by normal sedimentary fills. The Tarawan Formation is completely absent at Saint Paul and increased slowly in thickness southward. These variations reflect changes in sedimentation rate and tectonic disturbance (Fig. 13). 7. During Thanetian–Ypresian, a worldwide high stand of sea level causing major marine transgression. These conditions favored deposition of a thick monotonous succession of greyish and green shales (Esna Formation), being calcareous and dolomitic southward, with varying amounts of planktonic and benthonic forams. It seems that, the environment of deposition implies a continuation of the same conditions that prevailed during the Maastrichtian–Danian. The basin depocenter of the Esna Formation was shifted southwards around North Wadi Qena. This basin illustrates rapid lateral and vertical variations in thickness and bottom relief (Fig. 13). This kind of sedimentation represents the closing sedimentation interval for the studied sequence. The characteristic rock types; ortho- and allo-chemical components; sedimentary structures and other sedimentary features of the inferred facies are represented in Fig. 14. The different facies belts are arranged nearly in east–west oriented parallel belts and intersect the western shoulder of the Gulf of Suez. This pattern of sedimentation as well as facies changes may be chiefly controlled by the paleotopography that was inherited from K/P tectonism, which – in turn – resulted from the rejuvenation of the SAS. 7. The impact of tectonism on sedimentation The northern continental passive margin of Africa occupies the southern Mediterranean basin. In Egypt, this part of the margin was studied by Salem (1976), Meshref (1982) and Orwig (1982) besides others. Said (1990) believed that the geological history of the Egyptian Paleogene was dominated by tectonic events, which continued steadily or episodically into the Paleogene from Late Cretaceous tectonism. Strougo (1986) suggested a tectonic event activity (Velascoensis event) affected the depositional setting of many areas in south Egypt during Late Paleocene. Bandel and Kuss (1987) and

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Kuss (1989) pointed to Paleocene/Early Eocene uplift and related it to the later graben forming processes of the Red Sea/Gulf of Suez. Abu Khadra et al. (1994) recorded syn-tectonic Early and Middle Eocene event in the Southern Galala at Saint Paul and suggested that the central part of the Gulf of Suez was tectonically active site during Eocene. Integrated sedimentological and biostratigraphical studies resulted in recognition of four remarkable tectonic events (Fig. 15). These events are indicated mainly by:  The partial missing of the upper part of the Dakhla Formation (Fig. 7) and complete absence of some rock units at certain localities (e.g. Tarawan and Esna formations at Saint Paul locality).  The absence of some planktonic foraminiferal zones at different stratigraphic levels.  The major unconformity surface at the contact between Matulla and Sudr, hiatus between Sudr and Dakhla and between Dakhla and Tarawan formations.  Pronounced changes ((Fig. 7) in the style of sedimentation regimes (from chalk of Sudr to shales of Dakhla; from shales of Dakhla to carbonates of Tarawan and finally from carbonates of Tarawan to shales of Esna).  Vertical as well as lateral variations in the thickness of the same rock unit from one locality to another in areas geographically nearby each other (Fig. 13 and Table 2). 7.1. S/C tectonic event During the S/C, the studied rocks were affected by the main severe second phase of the SAS. Uplifting, marine regression and intensive erosion, formed a major unconformity throughout the study area. Furthermore, the normal slip of the northern bounding fault of the Southern Galala plateau might have been inverted into a reverse fault, as well as reactivating the other east–west oriented faults in the study area (Hussein and Abd-Allah, 2001). Uplifting of NED and Sinai took place during Coniacian and is related to the initiation of the Syrian Arc movement (Kerdany and Cherif, 1990). The second Syrian Arc phase was reported in several parts of Egypt (Moustafa and Khalil, 1995). This phase formed the fold/thrust belt of the SAS, which extended from Turkey to northern Egypt and Libya. This event took place at the basal part of Sudr Formation. It is traced at Saint Paul, Wadi El Dakhal and Gebel Millaha and documented by the absence of the Globotruncanita elevata Zone (Fig. 15). On outcrop, disconformity surface with pronounced relief cuts erosively into marls and shales of the Matulla Formation and is covered by caliche horizon. The base of the Sudr Formation includes glauconitic calcareous sandstone, which rests on irregular surface. Under this surface, the top of the Matulla Formation is tectonically deformed, stained with red color and interbedded with gypsum veinlets (Plate A1). Tectonically, the floor of the relatively shallow marine sequence of the Matulla Formation reconfigured to open marine continental slope and basin plain of the Sudr Formation. Furthermore, the regional pattern of sedimentation changed suddenly from siliciclastic shales to carbonates (Fig. 9). The time duration of this event is constant all over the area (nearly 3.8 million years, Fig. 15). 7.2. C/M tectonic event The C/M rocks might have been deposited in low areas flanking uplifted blocks and may have resulted from folds and faults in NED. The chalks of the Sudr Formation are deposited in the deeper shelf of an open marine environment. The Syrian Arc renewed tectonic activity resulted in the lowering of the northern outskirts of the

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Fig. 15. Conceptual, NNE/SSW oriented stratigraphic cross-section, reconstructed at different geologic times, showing the lateral and vertical lithofacies distribution at each time unit, as well as the effect of the different tectonic movements. Biozones and age assignment are based on Bralower et al. (1995) for Late Cretaceous, and Berggren and Pearson (2005) for Paleogene.

mid-Cretaceous shelf, where deeper shelf chalks and marl were deposited (Kuss, 1992). The paleogeographical reconstruction of NED indicates that the Southern Galala plateau and the down thrown block of the Northern Galala fault were the only depositional sites for the Campanian–Maastrichtian chalks (Hussein and Abd-Allah, 2001). In this study, the chalks of the Sudr Formation show slumping structures (Plate B4), this may be related to the third phase of the Syrian Arc movement (Moustafa et al., 1998). Toward the end of the Cretaceous, SAS were accompanied by magmatic activity (Hashad, 1980). This event took place at C/M boundary within the Sudr Formation and extends all over the area except North Wadi Qena site (Fig. 15). Its time duration is constant (nearly 4.2 million years). It is indicated by the complete absence of G. calcarata, G. subcarinata and G. aegyptiaca zones. Due to basinal nature of the sediments, the lithological impacts were not clearly imprinted on the sediments. The presence of this event during deposition of the Sudr Formation indicates clearly that the depositional basin of the Sudr Formation did not enjoy quiescence in spite of the apparent uniformity in sedimentation regime. 7.3. K/P boundary tectonic event A hiatus across the K/P boundary is evident all over the Egyptian territory (Said, 1990). On contrary, Scheibner et al. (2001) claimed that Saint Paul is considered one of the most complete sections across the K/P boundary in Egypt. Lüning et al. (1998) documented a complete section ranging across the K/P boundary on eastern Sinai. Together with the nowadays eroded Wadi Araba, the Southern Galala represents an area affected by the SAS and hence is a part of Said’s (1962) unstable shelf. In Southern Galala, compressional deformation resulted in uplift, erosion and non-deposition during K/P boundary. Bosworth et al. (1999) demonstrated that the SAS also affected regions at least as far south as the southern part of the Gulf of Suez (Esh El Mellaha Range and Gebel Zeit) due to far-field compressional stress. In contrast to the other events described before, this event has the most complex and inhomogeneous tectonic and depositional

history (Fig. 15). It could be subdivided into areas with long-range hiatuses across the K/P boundary (e.g. Gebel Millaha); and others with continuous and short-range ones (e.g. from North Wadi Qena to Saint Paul). The reason for these different depositional histories may lie in the fact that these areas were differently affected by SAS. Another possibility is the presence of a local paleohigh with nondeposition (there is no other evidence yet). A hiatus is present at the K/P boundary in most Egypt, which probably resulted from a combination of low sedimentation rates, reworking, changing circulation patterns and sea level changes (Lüning et al., 1998). This movement is coeval with K/P boundary (Haq et al., 1988), and could be traced on the top of Sudr Formation (contact with Dakhla Formation). The time duration of this event shows remarkable lateral and vertical variations (Fig. 15). This time ranges from nearly 5.3 million years around Gebel Millaha to 2.0 million years around other localities. The K/P boundary tectonic event is documented by absence of Late Maastrichtian P. palpebra and P. hantkenenoides zones and Early to Middle Danian P0, Pa, P1a, P1b, P1c zones. Sedimentologically, K/P event is recognized as submarine hardground populated by vertical to horizontal Thalassinoid tubes. Phosphate grains are scattered above the bioturbated surface. So, the basal part of the Dakhla Formation includes phosphatic material and reworked lithoclasts forming phosphatic packstone to pack/grainstone facies. The effect of this event on sedimentation is manifested by the change of the sedimentation regime from carbonates of the Sudr Formation to fine siliciclastic shales of the Dakhla Formation. On contrary, the paleo-relief of the sedimentary basin of the Dakhla Formation follows closely that of the Sudr Formation (Fig. 13) and the same open marine conditions, under which the Sudr Formation was accumulated, were still steady during the deposition of the Dakhla Formation. 7.4. Middle Paleocene–Early Eocene tectonic event Beginning with the Middle Paleocene time, the northern part of the study area, was subjected to a fourth phase of the Syrian Arc movements (Hussein and Abd-Allah, 2001). Strougo (1986) reports a Late Paleocene tectonic event that is marked by changes of

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depositional setting in several parts of Egypt. The Middle-Upper Paleocene marls and shales of the Dakhla Formation overlie unconformably the Maastrichtian chalks in the Southern Galala plateau. By the effect of this event, most of the Dakhla Formation, Tarawan and Esna formations are completely absent at Saint Paul and Wadi El Dakhal, where this movement reached its climax (nearly 6.6 million years, Fig. 15). Moving southwards, the echo of this movement is represented by short-ranging hiatuses within single foraminiferal biozones, which have very limited and restricted paleontological impact. It is believed here that the rejuvenation of NG/WAU associated with differential block faulting is the main reason for interpreting such events. Moustafa and Khalil (1995) stated that the age of folding in Wadi Araba and the South Galala plateau is certainly pre-Middle Eocene and most probably Late Cretaceous (early Late Senonian and later). Evidence for early Late Senonian folding is provided by the angular unconformity between the Lower and Upper Senonian rocks, and evidence for later rejuvenation is that Paleocene and Lower Eocene rocks are missing in the northern part of the South Galala plateau (Awad and Abdallah, 1966; Mazhar et al., 1979), whereas a complete stratigraphic section is found in its southern cliffs (south of latitude 28°450 N). In these southern cliffs, a complete Upper Cretaceous section is covered by Paleocene and Lower-Middle Eocene (Mazhar et al., 1979). This perhaps indicates that Wadi Araba structure was elevated high above the levels of the Paleocene and Early Eocene seas. A similar setting is also found in the southerneast and eastern scarp of Gebel Ataqa, where Upper Cretaceous rocks are unconformably covered by Middle Eocene limestones probably indicating Late Cretaceous folding (Ismail and Abdallah, 1966). The development of the SAS in the area under investigation occurred in four tectonic phases from Late Cretaceous to Early Eocene. Moreover, Sestini (1984) indicated that the Syrian Arc deformations continued up to Middle Eocene in other parts of Egypt. These tectonic phases are contemporaneous with the development stages of the Neotethyan Sea in the study area and other parts of northern Egypt. Several Cretaceous-Paleogene rocks that were deposited on the southern shores of the Neo-Tethyan Ocean contain significant hydrocarbon resources. Significant examples include Kerogen-rich limestones, which were developed in intrashelf basins, such as Late Aptian–Cenomanian of the Sinai ramp (Kim et al., 1999) or the productive Cenomanian–Turonian deposits of the Arabian Gulf (Alsharhan and Nairn, 1994). Such type of rocks is important reservoirs with porosities due to the interplay of different diagenetic alterations, or they are source rocks formed mainly during marine transgression. Similarly, K/P rocks are not an exception. They contain source-rock facies that are characterized by laminated argillaceous limestones and shales. If the down-faulted Galala-type sediments in the Gulf of Suez are similar, we may expect to find mature organic-rich carbonate and shale deposits there that were formed in intrashelf basins of Late Campanian–Early Eocene. For example, the source rock of the Zafarana field in the Gulf of Suez is probably composed of organic-rich Senonian–Eocene carbonate deposits (Kuss et al., 1999).

8. Summary and conclusions The K–P sequence rests unconformably over the fine siliciclastic shales of the Matulla Formation (Santonian) and is covered conformably to unconformably by the limestone of the Thebes Formation (Late Ypresian). The studied sequence is about 120 m thick. Rock units are from base to top: Sudr Formation (Campanian– Maastrichtian); Dakhla Formation (Danian–Selandian); Tarawan Formation (Selandian–Thanetian) and Esna Formation (Thanetian–Ypresian).

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18 planktonic foraminiferal zones have been defined covering the K–P sequences under study. These zones are: Globotruncana ventricosa Zone for the Campanian age; Gansserina gansseri, Contusotruncana contusa, Recimguembelina fructicosa, Pseudohastigerina hariaensis, P. palpebra and Plummerita hantkenenoides zones for the Maastrichtian age; Praemurica incostans, P. uncinata, Morozovella angulata and Praemurica carinata/Igorina albeari zones for the Danian age; Igorina albeari, Globanomanlina pseudomenradii/Parasubbotina variospira, Acarinina subsphaerica, Acarinina soldadoensis/Globanomanlina pseudomenardii and Morozovella velascoensis zones for the Selandian–Thantian Age and Acarinina sibaiyaensis, Pseudohastigerina wilcoxensis/Morozovella velascoensis zones for the earliest Ypresian Age. Several hiatuses interrupting these planktonic foraminiferal zones are repeated vertically at certain levels. Such zonal interruption is reflected in sedimentation regime during the K–P (Campanian–Ypressian) time. Integration of detailed field investigations and microfacies analysis revealed the predominance of four sedimentary facies belts forming a south-southwest gently-dipping slope to basin transect. The arrangements of these facies also take the form of east–west oriented parallel belts. This pattern of sedimentation was controlled chiefly by the paleorelief inherited from the K–P tectonism. Open marine incursion was linked with the initiation of a deep subsiding elongated trough within the continental shelf during the Maastrichtian. As a result of basin floor subsidence and relative sea level rise, thick monotonous pelagic carbonate sedimentation (Sudr Formation) of open marine environments prevailed during the Maastrichtian time. At the end of the Maastrichtian, a short regressive phase took place, producing a break in sedimentation (K/P boundary). Afterwards, a major marine transgression took place during the Early and Middle Paleocene, which favored deposition of a thick monotonous succession of greyish and green shales (Dakhla Formation). With the passage of time, the basin started to lose its morphologic expression and reflects the evolution from an old distally steepened ramp to a new homoclinal ramp. Upon the latter ramp, sedimentation of the Tarawan Formation began in an open outer shelf domain. A worldwide high stand of sea level causing major marine transgression during P/E transition, led to deposition of a thick monotonous succession of greyish green shales (Esna Formation). This kind of shale deposition represents the closing sedimentation interval in the studied localities. Integration of field criteria, biostratigraphical and sedimentological investigations led to recognition of four tectonic events, which took place on the NED during the K–P time. These events are the S/C boundary; the C/M boundary; the K/P boundary and the Middle Paleocene/Early Eocene boundary. Acknowledgements The authors are indebted to late Prof. Ahmed S. Kassab, Geology Department, Assiut University for field assistance, guidance and encouragement. Prof. Dr. W. Bishara is thanked for improving English spelling and grammar. The authors are grateful to the editorial board and anonymous referees for valuable comments which have greatly improved this paper. References Abdallah, A.M., Abu Khadrah, A.M., Darwish, M., Helba, A.A., 1996. Lithostratigraphy of Gebel Safariat Area, SW Sinai, Egypt. In: 3rd Int. Conf. Geol. Arab World, Cairo Univ., pp. 455–478. Abdel-Kireem, M.R., Samir, A., 1995. Biostratigraphic implications of the Maastrichtian-Lower Eocene sequence at the North Gunna section, Farafra Oasis, Western Desert, Egypt. Mar. Micropaleontol. 26, 329–340. Abu Khadra, A.M., Darwish A.M., El Azabi, M.H., 1994. Contribution to the faulteddown Eocene limestones of the Southern Galala plateau (St. Paul area), Gulf of Suez, Egypt. In: 2nd Int. Conf. on the Geol. of the Arab World, Cairo Univ., pp. 418–437.

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