Sequence stratigraphy of the Upper Cretaceous

of central-east Sinai, Egypt

1*S. LuÈ ning, {A. M. Marzouk, {A. M. Morsi and *J. Kuss

* University of Bremen, FB5-Geosciences, PO Box 330440, 28334, Bremen, Germany; 1present address:Lasmo plc, 101 Bishopsgate, London, EC2M 3XH, UK{Tanta University, Faculty of Science, Geology Dept., Tanta, 31511, Egypt{Ain Shams University, Faculty of Science, Geology Dept., Abbassia, Cairo, Egypt

Revised manuscript accepted 6 November 1997

Deposition of the Upper Cretaceous of central-east Sinai was controlled by a long-term transgressivephase and several higher order sea-level ¯uctuations. The paper gives a ®rst sequence stratigraphicinterpretation for this interval in the region, based on detailed sedimentological, biostratigraphical andpalaeoecological investigations of 13 Turonian-Maastrichtian sections and a review of all publisheddata. Six main facies zones have been differentiated: coastal mud ¯ats with tidal channels, gypsiferoussabkha plains, peritidal siliciclastics, peritidal carbonates, high- and low-energy carbonate inner shelffacies, and microfossil-rich outer shelf pelites. The biostratigraphy is based mainly on planktonic fora-minifera, calcareous nannofossils, ostracods and ammonites. The study is restricted to an area thathad been tectonically rather quiet during the Late Cretaceous-early Tertiary, lying south of the SyrianArc intraplate foldbelt which experienced major uplifting during this period. Within the Turonian toMaastrichtian interval, six major sequence boundaries have been recognized. Cycle duration variedbetween 4 and 9 Ma, which indicates a cycle order intermediate between 3rd and 2nd. Correlationwith other sea-level reconstructions for the region (Egypt, Israel, Jordan, Tunisia) points to a more orless synchronous regional sea-level development. Comparison of the regional sequences with the`eustatic' model of Haq et al. (1987) involves uncertainties; nevertheless, some of the sea-level ¯uctu-ations recorded in Sinai may be related to worldwide eustatic sea-level changes.

# 1998 Academic Press

KEY WORDS: Sinai; Egypt; Cretaceous; biostratigraphy; sequence stratigraphy.

1. Introduction

Deposition of the Upper Cretaceous of Sinai and of the Negev Desert was

controlled by a long-term transgressive phase and several higher order sea-level

¯uctuations with depositional environments ranging from terrestrial to

hemipelagic. The depositional history of Sinai and of the Negev Desert during

this period has been reported previously by several authors (e.g., Bartov et al.,1972; Flexer et al., 1986; Cherif et al., 1989a, b; Lewy, 1990; Kuss, 1992; Orabi,

1992, 1993; Ziko et al., 1993; Kora & Genedi, 1995). Nevertheless, a

comprehensive sequence stratigraphic model for these successions has not been

published. This study, therefore, aims at giving the ®rst sequence stratigraphic

interpretation for the Upper Cretaceous of central-east Sinai on the basis of

detailed sedimentological, biostratigraphical, and palaeoecological investigations

of 13 Turonian-Maastrichtian sections (Figure 1). The available literature on this

interval in central-east and central-west Sinai has been reviewed and included in

the sequence model. The study area is situated in a region that was tectonically

rather quiet during the Late Cretaceous-early Tertiary, lying south of the Syrian

Arc intraplate fold belt which experienced major uplift during this period. The

sea-level changes described in this study are, therefore, expected to possess at

Cretaceous Research (1998) 19, 153±196 Article No. cr971014

0195±6671/98/020153 + 44 $30.00/0 # 1998 Academic Press

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154 S. LuÈning et al.

least regional character and may be largely independent of local compressional

processes in the fold belt. However, larger-scale loading and relaxation effects in

terms of a foreland basin model cannot be fully excluded.

2. Regional setting

During the Late Cretaceous, Sinai was part of the broad northern shelf of the

Afro-Arabian Plate. The shelf margin is inferred to have been located near the

present-day northern coastline of Sinai (Bein & Gvirtzman, 1977; Ginzburg &

Gvirtzman, 1979). The Gulf of Suez and Gulf of Aqaba rifts, which bound the

Sinai microplate today, were still closed. While northern Sinai and the adjacent

Negev were affected by transpressive movements from the Late Turonian

onwards, southern and central Sinai (the study area) are believed to have

remained tectonically rather quiet throughout the Mesozoic and Early Tertiary

(Said, 1962; Cohen et al., 1990; Kerdany & Cherif, 1990). The SE-vergent,

NE-SW trending domal anticlines in northern Sinai are part of the `Syrian Arc'

(Krenkel, 1924, 1925) which represents an intraplate fold belt extending from

Egypt to Syria, and was formed by Late Cretaceous to Recent inversion of Late

Triassic/Liassic half-grabens (Moustafa & Khalil, 1990; Chaimov et al., 1992;

Shahar, 1994). Following the nomenclature of Said (1962), the inversion zone is

also termed `stable shelf ', and the southern tectonically inactive block is termed

`unstable shelf ' (Figure 1). With one exception, the sections studied here are

situated on the stable shelf.

Owing to the intermediate position between mainland Egypt and Israel, two

different lithostratigraphic systems have been used for the Upper Cretaceous in

Sinai. While the Egyptian formations are mainly based on Ghorab (1961), the

Israeli nomenclature was developed by Flexer (1968) and Bartov et al. (1972).

Correlations of the different schemes (Figure 2) are included, for example, in

Figure 2. Correlation of Egyptian and Israeli formations in eastern Sinai (modi®ed after Kora &Genedi, 1995).

Sequence stratigraphy in Egypt 155

Kerdany & Cherif (1990), Kuss (1992) and Kora & Genedi (1995). In the

following section, a short description of the different formations in the study is

given; detailed lithological descriptions and remarks on differentiation of the

formations in central-east Sinai can be found in Bartov et al. (1972), Ziko et al.(1993) and Kora & Genedi (1995). The Abu Qada Formation (=Ora Shales in

the Israeli nomenclature) consists of mainly pelitic lithologies and is attributed to

the late Cenomanian to early Turonian. These soft lithologies are overlain by

hard limestones of the upper Turonian Wata Formation (=Gero®t Formation).

The Coniacian-Santonian succession is included in the Matulla Formation

(=Zihor Formation and lower part of Sayyarim Formation), consisting mainly of

calcareous siliciclastics, carbonates and inner shelf shales. Towards the north, the

siliciclastics are replaced by carbonates which are grouped into the Themed

Formation in this area. These units are overlain by the Campanian-Maastrichtian

Sudr Chalk (upper Sayyarim and Ghareb Formations) which is composed of

chalks, cherts and limestones. Isopach maps for the different Upper Cretaceous

formations can be found in Bartov & Steinitz (1977). Upper Cretaceous deposits

play a signi®cant role in hydrocarbon plays in the petroliferous Gulf of Suez Basin

(e.g., Sestini, 1995) as well as in the rare hydrocarbon discoveries in northern

Sinai (see Alsharhan & Salah, 1996).

3. Materials and methods

During two ®eld expeditions in 1995 and 1996, twelve Upper Cretaceous

sections were measured in central-east Sinai and one section on the Gebel Areif

El Naqa anticline in northeast Sinai (Figure 1). During the ®eldwork, special

attention was paid to sedimentary structures and horizons showing clear

lithological changes. For thin section analysis, 48 hand specimens from

limestones and sandstones were taken. For microfossil investigations, chalks,

marls, and shales were sampled at intervals of 1±3 m. A total of 171 samples was

collected. For extraction of foraminifera and ostracoda the pelitic samples were

washed twice using a 63 mm sieve after treatment with H2O2 and a highly

concentrated tenside ('REWOQUAT'), respectively. The microfossil residue was

then fractionated into four grain size classes for easier handling. The planktonic

foraminifera were identi®ed under the light microscope (by LuÈning). For the

calcareous nannoplankton, smear slides were prepared using techniques described

in Bramlette & Sullivan (1961) and Hay (1961, 1965). The slides were examined

(by Marzouk) at a magni®cation of about �1250 by both cross-polarized and

phase-contrast illumination. Ostracods were picked from all samples and were

investigated by Morsi.

4. Biostratigraphy

The biostratigraphy in this study is based mainly on planktonic foraminifera and

calcareous nannofossils; macrofossils yielded only a little supplementary data (see

below). For palaeoecological reasons, biostratigraphy using planktonic

foraminifera and calcareous nannofossils is restricted to the more hemipelagic

parts of the succession. The facies of the upper Coniacian, Santonian and

Maastrichtian are often well suited for biostratigraphic studies, whereas the

shallow-marine deposits of the Turonian to middle Coniacian in the study area

often do not contain biostratigraphically indicative species. Although intense

156 S. LuÈning et al.

silici®cation restricts microfossil preservation to several intervals of the mostly

hemipelagic Campanian, planktonic foraminiferal faunas of varying abundance

have been found in the unsilici®ed parts of the succession.

The Late Cretaceous hemipelagites of Sinai have been previously studied

biostratigraphically on the basis of planktonic foraminifera by several authors.

While some workers focused on speci®c time intervals (Cenomanian-Turonian:

e.g., Cherif et al., 1989a; Orabi, 1992; Coniacian-Santonian: Ismail, 1993;

Maastrichtian-Eocene: e.g., Said & Kenawy, 1956; Abdelmalik et al., 1978;

Cherif et al., 1989b, Shahin, 1992), others considered the whole of the Late

Cretaceous interval (Shahin & Kora, 1991; Hewaidy et al., 1991; El Sheikh,

1995; Ayyad et al., 1996). In this contribution we use the general Tethyan

biozonation scheme for planktonic foraminifera by Robaszynski et al. (1984), and

for the Turonian to Lower Coniacian, the scheme of Caron (1985) is adopted.

However, some regional peculiarities exist, namely the absence of Globotruncanitacalcarata in central-east Sinai, which is also supported by the study of Shahin &

Kora (1991). This may in part be related to silici®ed-phosphoritic lithologies in

the uppermost Campanian (Taba section, Figures 7, 8) from which no planktonic

foraminifera could be recovered. Findings of G. calcarata were previously

reported from northeastern Sinai (Hewaidy et al., 1991) and Israel (Almogi-Labin

et al., 1986; Gvirtzman et al., 1989), whereas in other sections in Israel (e.g.,

Almogi-Labin et al., 1993) and western Sinai (Cherif et al., 1989b; Orabi, 1991)

this species is missing. The ®rst occurrence of G. aegyptiaca in the G. calcarataZone was previously reported by Almogi-Labin et al. (1986) from Israel. It has

been found at the base of the G. falsostuarti Zone in northeastern Sinai (Hewaidy

et al., 1991). This contrasts with the Tethyan scheme of Robaszynski et al. (1984)

who postulated a ®rst occurrence within the lower Maastrichtian, representing the

base of the G. aegyptiaca Zone of Caron (1985). Owing to the general absence of

G. calcarata and the early occurrence of G. aegyptiaca, the base of the G.falsostuarti Zone, and therefore the Campanian-Maastrichtian boundary (see also

general boundary discussion in Odin, 1996), remains rather poorly de®ned in the

Taba section (section M, Figure 8A). In this section, the interval directly below

the G. gansseri Zone is attributed to the lower Maastrichtian G. falsostuarti Zone

(as opposed to the upper Campanian G. calcarata Zone) because of the absence

of G. subspinosa. Some characteristic planktonic foraminifera from the Upper

Cretaceous of eastern Sinai are illustrated in Figure 3.

Biozonation schemes for benthonic foraminifera in the Upper Cretaceous in

Sinai and the Gulf of Suez were presented by Ansary & Tew®k (1969), Shahin &

Kora (1991), Hewaidy et al. (1991), and Hewaidy & El Ashwah (1993).

However, because the resolution is very poor and the occurrence of index species

is considered to depend strongly on facies, benthonic foraminifera have not been

utilized for biostratigraphic purposes.

Only a few studies of calcareous nannofossils from the Upper Cretaceous of

Sinai have been published (Arafa & El Ashwah, 1988; Arafa, 1991; El Sheikh,

1995; Swedan et al., in press). More extensive investigations have been carried

out in neighbouring Israel (e.g., Eshet & Moshkovitz, 1995). Biozonation of

calcareous nannofossils herein is based on the concepts of Crux (1982) for the

Turonian to early Campanian, and of Eshet & Moshkovitz (1995) and Perch-

Nielsen (1979) for the late Campanian to Maastrichtian. Correlation of

calcareous nannofossil biozones with those of planktonic foraminifera shows

varying degrees of correspondence with the schemes published by Bralower et al.

Sequence stratigraphy in Egypt 157

Figure 3. Planktonic foraminifera from the Upper Cretaceous of eastern Sinai (sections A5, A6,A7, C, D; sections A5 and A7 near A6); scale bars � 100 �m. 1, 2, Abathomphalus mayaroensis,late Maastrichtian, specimen 1 from sample A7-11, specimen 2 from sample A7-15; 3, Plummer-ita reicheli (sensu Masters, 1993), latest Maastrichtian, both specimens from sample A5-17; 4, 5,6, Globotruncana aegyptiaca, late Maastrichtian, specimen 4 from sample C1-5, specimen 5 fromsample C1-6, specimen 6 from sample C1-7. 7, 8, Dicarinella asymetrica, Santonian, specimen 7from sample A6-14, specimen 8 from sample A6-8; 9, Whiteinella archaeocretacea, Santonian,sample A6-5; 10, 11, Dicarinella concavata, Coniacian-Santonian, both specimens from sampleD1-3; 12, 13, Marginotruncana sinuosa, Coniacian-Santonian, both specimens from sample D1-3.

158 S. LuÈning et al.

(1995) for the Tethys, and Eshet & Moshkovitz (1995) for Israel. Although a

good correlation with these schemes has been observed for the late Coniacian to

Campanian, the Maastrichtian correlation comes closer to the schemes in Bolli

et al. (1985). The main differences in the correlation of our data from the Taba

section with the schemes in Bralower et al. (1995), Eshet & Moshkovitz (1995)

and Gvirtzman et al. (1989) are the early ®rst occurrences of Lithraphiditesquadratus near the base, and Micula murus in the upper Gansserina gansseri Zone.

The differences may be explained, at least in part, by the strong palaeoecological

susceptibiliy of the two species. While L. quadratus is interpreted to be

characteristic of low productivity conditions, M. murus has been found to have

bloomed in high productivity settings. Some characteristic calcareous nannofossils

from the Upper Cretaceous of eastern Sinai are illustrated in Figure 4.

Detailed biozonation schemes based on ostracoda have been proposed for the

Cenomanian-Turonian (Rosenfeld & Raab, 1974) and Coniacian-Maastrichtian

(Honigstein, 1984) in Israel. Because some important marker species have not

been recorded in our material, biozonation by ostracods in the sections studied is

of lower resolution than in the schemes from Israel, although this group

occasionally provides valuable information.

For the Turonian-Coniacian in this study, supplementary biostratigraphic data

are derived from echinoids (Hemiaster solignaci), oysters (Gyrostrea antwani), and

ammonites (Choffaticeras sp.). Another ammonite species has been found around

the Campanian-Mastrichtian boundary (Solenoceras humei). During ®eldwork, few

these fossils were found (Figure 14A-C). However, other studies in which a

focused search for ammonites in the ®eld had been carried out, showed that these

fossils may provide a valuable biostratigraphic tool in Sinai (e.g., Lewy, 1975;

Lewy & Raab, 1976; Bartov et al., 1980; Abdel-Gawad, 1990; Abdel Gawad etal., 1992). It must be noted that the oyster Pycnodonte vesiculare has been found

in the Upper Coniacian/Lower Santonian in the Taba section, which falls well

within the Upper Coniacian-mid Maastrichtian stratigraphic range indicated for

this species by Malchus (1990). P. vesiculare does not, therefore, represent a

stratigraphic marker for the Campanian-Maastrichtian only, as misinterpreted by

several authors (e.g., Refaat, 1993; Orabi & Ramadan, 1995). Recent

investigations of authigenic K-Ar systems of silicate fractions from Upper

Cretaceous rocks in Israel have yielded K-Ar ages similar to the suggested ages of

deposition (Steinitz et al., 1995), providing an independent geochronological

method for age determination in the rock successions studied.

5. Facies

In this study, six facies zones are differentiated in the Upper Cretaceous of

central-east Sinai. The main depositional elements include coastal mud ¯ats with

tidal channels, gypsiferous sabkha plains, peritidal siliciclastics, peritidal

carbonates, high and low energy carbonate inner shelf facies, and microfossil-rich

outer shelf pelites. An idealized block diagram summarizing all facies developed

during the Late Cretaceous in the study area is given in Figure 5. The diagram is

intended as a reference for the interpretation of vertical facies changes in terms of

Walther's law of succession of facies, and is used to illustrate potential positions

of inter®ngering facies (see descriptions below). This facies model has to be

regarded as a predictive tool and not as a palaeogeographic reconstruction for

a certain time interval; hence, no scale or orientation is given. Detailed

Sequence stratigraphy in Egypt 159

palaeogeographic maps of central Sinai within a sequence stratigraphic framework

are presented and discussed further below. Concerning the interpretation of the

facies diagram it is important to note that, for any given time interval, central-east

Sinai was dominated by only one to a maximum of three main facies (see below),

suggesting that the study area represents only a small part of a much larger

continent-basin pro®le.

In the following sections, the different facies are described and an

environmental interpretation is given. Facies investigations in central-east Sinai

Figure 4. Calcareous nannofossils from the Upper Cretaceous of eastern Sinai (sections A1, A6,C, M, T2). Scale bar � 15 �m. 1, Lucianorhabdus cayeuxii, Campanian; sample A6-18; 2, Watz-naueria barnesae, late Maastrichtian, sample M1-30; 3, 6, Aspidolithus parcus constrictus, earlyCampanian, sample A6-22; 4, 19, Eiffellithus eximius, late Santonian, sample A6-14; 5, 8, Aspido-lithus parcus expansus, early Campanian, sample A6-22; 7, 10, Quadrum sissinghii, late Campanian,sample M1-22; 9, Tetrapodorhabdus decorus, late Maastrichtian, sample C1-3; 11, Lucianorhabdusarcuatus, late Santonian, sample A1-30; 12, 16, Reinhardtites anthophorus, around Campanian-Maastrichtian boundary, sample A6-21; 13, 18, Calculites obscurus, late Santonian, sample A1-31;14, Quadrum tri®dum, late Campanian, M1-23; 15, 17, 22, Zeugrhabdotus pseudanthophorus,around Santonian-Campanian boundary, sample A6-17; 20, Eiffellithus turriseiffelii, Campanian,sample T2-35; 21, Reinhardtites levis, around Campanian-Maastrichtian boundary, sample A6-21.

160 S. LuÈning et al.

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Sequence stratigraphy in Egypt 161

have been carried out previously by Bartov et al. (1972), Lewy (1975), Orabi

(1992, 1993), Ziko et al. (1993), Khalifa & Eid (1995), Kora & Genedi (1995),

Ahmed (1995); their results, as well as data from neighbouring regions, have been

included in our model.

5.1. Supra- and intertidal mud¯ats, sabkhaDescription. This facies is mainly developed in the Lower Turonian (e.g., section

B, Figure 6) and Coniacian-Santonian but also forms thin units in the Upper

Turonian and Upper Campanian parts of the succession. These are characterized

by alternating, partly ®nely laminated, mottled reddish-greenish soft siltstones,

shales (Figure 10A) and, in the Lower Turonian, primary gypsum. In many cases,

the siltstones are cross-bedded with occasional intercalations of channel ®lls of

®ne sandstone. In the Coniacian/Santonian of Taba (Figure 7), reddish-brown

clay horizons are developed. In the phosphoritic Upper Campanian of the same

section (Figure 8), a marl with abundant plant fragments (Figure 10C) and

lacking foraminifera has been found.

Interpretation. The ®ne-grained siliciclastic sediments are interpreted to have been

deposited in silty/shaly supra- to intertidal mud¯ats located in deltaic and tidal

interdistributary areas which have been sourced by overbank and crevassing

processes. Lamination, ripples and silt-shale alternations re¯ect episodic ¯ooding

events of varying intensities. Intercalated ®ne sandstones (e.g., in the Lower

Turonian of section B and the Coniacian-Santonian of section M) represent

deltaic and tidal distributary or crevasse channel deposits. The occurrence of

gypsum is restricted to a sabkha setting developed mainly during the early

Turonian which is related to a pronounced arid climate at this time (section B)

and, although only in a few layers, during the Coniacian-Santonian (section M).

Red-green layering and mottling re¯ect strong vertical and lateral changes in the

diagenetic redox potential, related to bioturbation, hydrological and climatic

conditions leading to a complex Fe2 + /Fe3 + distribution. Similar colour patterns

can be found in modern tidal ¯ats, for example, along the German North Sea

coast. The plant fragments in the Upper Campanian marl point to the

development of a vegetated swampy marsh facies during this period.

5.2. Siliciclastic delta and tidal sand ridgesDescription. Quartzose ®ne to medium sandstones of this facies are developed in

the Coniacian-Santonian section, and are intercalated with carbonates. The ®ne

sandstones are composed of ®ne, evenly laminated units alternating with dm-scale

erosive cross-bedded intervals (Figure 12E) and sharply bounded, intensely

slumped horizons. Slump fabric is characterized by asymmetric slump folds,

slump-fault planes and slump-brecciation on a cm-scale (Figure 12B-D). In the

Taba section (section M), pseudonodules of coarse to medium sandstone have

sunk into layered ®ne siltstone and deformed the underlying sediment

(Figure 12F). This process is likely to occur if a ®eld of coarser-grained sand

ripples migrates over unconsolidated ®ner-grained sediment which leads to

loading effects and pseudonodule formation (Allen, 1982). The medium

sandstones are arranged in tabular cross-beds of m-scale with vertically variable

dip directions, or are horizontally bedded. Locally, discrete layers rich in

Thalassinoides are developed. In the Taba section (section M), a clear coarsening

162 S. LuÈning et al.

upward trend from siltstone/®ne sandstone to medium sandstone can be

recognized within a 10 m interval (Figure 7).

Grain composition re¯ects a high sandstone maturity. Grains consist

predominantly of well to moderately sorted quartz and accessory feldspar and are

angular to sub-angular/sub-rounded (Figure 12G). All sandstones contain varying

amounts of calcareous, often dolomitic, matrix with lithological transitions to

sandy (dolo-) wackestones, the latter being characterized by the occurrence of

bivalves, oysters, green algae, ostracods, and gastropods (Figure 12I). The

inter®ngering of siliciclastic and carbonate facies is also documented in ®ne

alternations (mm-cm-scale) of rippled ®ne to medium sandstone and dolomitic

micrite (Figure 12H). In other examples, clusters of quartz grains accumulated in

a dolomitic matrix. At some places in the Santonian and Campanian (e.g., in

section T2), the sandstones have a phosphoritic character with the phosphorus

contained in the matrix or in the poorly sorted, angular to peloidal phosphoritic

components which also include teeth and bone fragments. Glauconitic grains are

often found. Porosity in all sandstones is relatively high.

In central-east Sinai, the Coniacian-Santonian siliciclastic interval was

previously studied by Khalifa & Eid (1995). The siliciclastics from SW Sinai and

the Gulf of Suez were comprehensively described by Refaat (1993) and Enani et

al. (1994) based on surface sections and offshore cores. Enani et al. (1994)

convincingly showed that the sandstones in SW Sinai were deposited as ENE-

oriented linear bar structures, several tens of km long, 5-10 km wide and 3-20 m

thick.

Interpretation. The detailed studies of the Coniacian-Santonian siliciclastics by

Enani et al. (1994) in western Sinai and the Gulf of Suez showed that the

sandstones in that area are arranged in ENE-oriented linear tidal bars with their

long axes perpendicular to the coastline. By analogy, a similar interpretation is

made for the Coniacian-Santonian sandstones in central-east Sinai. However,

orientations of the bars may be different because the coastline in central Sinai

during this period was strongly curved (Refaat, 1993; Said, 1990). In the study

area, bars are inferred in two sections (sections M and D; Figures 7, 9B). In the

Coniacian of the Taba section (section M), a typical shallowing-upward trend,

including the sandbar facies, is developed. The succession begins with a low

energy inner carbonate shelf facies, followed by ®nely laminated, partly ¯asered or

rippled siltstones and ®ne sandstones which are interpreted as deposits of the

lower shoreface or offshore transition zone. This assumption is based on the

vertical position of the unit within the succession, the typical sedimentary

structures, the absence of bioturbation pointing to reworking by (storm) waves,

and the slumped horizons which suggest the presence of a gradient. The overlying

tabular cross-bedded medium sandstones are attributed to tidal sandstone bars

sensu Enani et al. (1994) whereas the following cross-bedded sandy oyster

¯oatstones point to an inter-bar setting. The shallowing upward succession is

topped by a typical reddish silty mud ¯at facies dissected by small sandy

channels.

Proximity and inter®ngering of the tidal sandstone facies with the nearby inner

neritic carbonate facies is demonstrated by several features; namely, the abundant

calcareous matrix of the siliciclastics, lithologic transitions to sandy bioclastic

wackestones, and small-scale sandstone-micrite alternations. While the high

maturity of the sandstones with predominantly quartz grains points to

Sequence stratigraphy in Egypt 163

considerable transport on the African Craton, the often subangular grain-shapes

seem to exclude prolonged transport and indicate more regional source areas.

The phosphoritic character of the Santonian-Campanian sandstones is connected

with basinwide upwelling processes at the southern Tethyan margin during this

period.

5.3. Calcareous peritidal faciesDescription. In central-east Sinai, the peritidal facies has only been found in the

Upper Turonian. A typical example is developed in the Upper Turonian near

Themed (section N1) where several subtidal-intertidal shallowing-upward cycles

of dm-scale were observed (Figure 10B). While the subtidal lower parts of the

cycles are composed of strongly bioturbated (Thalassinoides) wackestones and

packstones containing miliolids, gastropods and oysters, the upper parts of the

cycles are characterized by micritic laminites with even or wavy lamination. In

some intervals, laminated stromatolites have been found which, according to

Monty (1967), are mainly developed in the intertidal zone. In other localities

(e.g., section B), inversely graded peloidal layers with fenestrate and birds-eye

fabrics alternate with dark micritic horizons containing angular micritic clasts

(Figure 10E, F).

Interpretation. The calcareous peritidal facies is only developed in settings with

little terrigenous input within the studied interval. Consequently, suitable

conditions only occurred during the late Turonian, because during the early

Turonian and Coniacian/Santonian, the peritidal zone was characterized by

siliciclastic mud ¯ats and, only during the Conaician/Santonian, by tidal bars.

5.4. Calcareous inner shelfÐlow energy faciesDescription. The calcareous, low energy inner shelf facies is mainly developed in

the Upper Turonian and, in the northern part of the study area, also in the

Coniacian. This facies consists of partly dolomitic wackestones, containing

miliolids (Figure 10I) (e.g., Quinqueloculina sp.) and other benthonic foraminifera,

ostracods, gastropods, green algae (Figure 10G), molluscan shells (including

oysters, Figure 10H), serpulids (Figure 10G), micritic clasts, peloids, echinoids

and bryozoa. Planktonic foraminifera are absent or very rare. Quantitative

distribution of the different components varies widely in the studied successions.

Observed end members include miliolid wackestones, ostracod wackestones,

wackestones with green algae and molluscan shells, wackestones with low-diverse

benthonic foraminifera and very few planktonic foraminifera, micritic oyster

¯oatstones, gastropod wackestones and ¯oatstones, and echinoid ¯oatstones.

Some of the low energy intervals are strongly bioturbated by Thalassinoidesburrows (Figure 11A). A Coniacian E-W-trending, low energy oncolite belt in the

Themed area has been described by Lewy (1972).

Interpretation. See higher energy calcareous inner shelf facies below.

5.5. Calcareous inner shelfÐhigher energy faciesDescription. The higher energy inner shelf facies can be found mainly in the Upper

Turonian and, in the northern part of the study area, also in the Coniacian

succession. This is similar to the stratigraphic distribution of the low energy

calcareous facies. The high energy facies is composed of well-sorted packstones

and grainstones with characteristic sedimentary structures such as dm-scale low-

164 S. LuÈning et al.

amplitude trough cross-strati®cation, wave and current ripples (Figure 11C, F),

normal grading, erosive bases, scours (Figure 11B), and sharply bounded

conglomeratic tempestites (dm-m scale) which are often composed of dolomitic

clasts. The composition and fabric of these clasts suggests that they were

reworked from the low energy facies (see above). The main components are

peloids, green algae (Figure 11E), gastropods, ostracods, echinoid spines, and

benthonic foraminifera (Figure 11H). The components are often surrounded by

one to a few micritic layers so that a super®cial oolitic fabric resulted

(Figure 11G). In some cases, intervals with multilayered ooliths have been found.

Typical microfacies types include oolitic grainstones with abundant green algae,

oolites with molluscan shells and gastropods, and green algae-gastropod

grainstones. Reworking is also testi®ed by abundant, large, angular micritic clasts

(Figure 11D, G) and micritic internal sediment in molluscan tests. Porosity of the

oolites is relatively high with cores leached in many cases. In the ?upper Turonian

of section N1, several horizons of reworked rudists have been found which

indicate the presence of rudist buildups.

In the Campanian and lower Maastrichtian deposits at some localities,

phosphatic higher energy deposits are developed, containing abundant

subrounded, structureless phosphoritic clasts and ®sh teeth. These rocks have a

dark brown to black weathered appearance in the ®eld and a characteristic smell

when crushed.

Interpretation. The calcareous inner shelf is interpreted to consist of a complex,

patchy system of higher energy shoals which accumulated material reworked from

low energy, micritic, inter-shoal areas. Such a relationship is suggested by the

clear evidence of reworking and facies inter®ngering presented above. During

redeposition in the shoals, the reworked components were often encrusted by thin

super®cial ooid coatings. On occasion, this process continued to produce

multilayered ooids, probably under localized high-energy conditions. Although

statistical, high resolution microfacies analyses have not been carried out in

central Sinai for the Upper Cretaceous, differentiation into several microfacies

end members suggests that a number of subenvironments may be recognized.

The low energy microfacies includes lagoons dominated by miliolids, ostracods,

green algae, gastropods, echinoids, smooth-shelled oysters or a combination of

these organisms. The composition of the shoals, however, may have been strongly

dependent on the type of the neighbouring low energy facies which provided most

of the material for them. Rudist banks and oyster bioherms (Bartov & Steinitz,

1982) represent higher energy subfacies with signi®cant autochthonous carbonate

production. Because the facies assemblage described is interpreted to possess a

relatively high potential for autocyclic migration processes (see also Powell et al.,1996), lateral reconstruction of carbonate subfacies as well as detailed sea-level

interpretations based on (non-peritidal) inner shelf carbonate facies distribution,

is complicated and was not, therefore, attempted in this study.

5.6. Hemipelagic outer shelf faciesDescription. The hemipelagic facies is widely developed from the Santonian to

Maastrichtian of central-east Sinai. Similar facies, although with a shallower

character, have been found around the Cenomanian-Turonian boundary (section

B; Figure 6) and in the Lower Coniacian (section M). The hemipelagic deposits

are dominated by grey and white chalks, soft greenish to grey marls and

Sequence stratigraphy in Egypt 165

Figure 6. Section B (N0 Sheikh Attiya). Legend in Figure 8C; for location, see Figure 1.

166 S. LuÈning et al.

calcareous marls containing varying amounts of planktonic foraminifera and

calcareous nannofossils, which in many cases can be easily isolated for

biostratigraphic investigations. Bioturbation is observed at many levels. Regular

alternations of chalky calcareous marls and softer marls on a cm/dm-scale are

discernible in some parts of the succession, especially in the Upper Maastrichtian

(section F; Figure 1).

Within the upper Coniacian to Campanian units, bedded ¯ints with thicknesses

ranging from dm to 10 m are frequently developed and in some cases alternate

with chalks and calcareous marls. Locally, the ¯int is ®nely laminated. The ¯int is

interpreted to consist of SiO2 derived from diatoms and radiolaria which today

are dissolved (Moshkovitz et al., 1983). In the Campanian ¯int succession of

Taba (section M, Figure 8A), some horizons are characterized by pronounced

swell structures with chert swells of dm-m scale protruding several dm into the

overlying carbonate series. Similar structures were previously described from the

Sinai/Negev area as `chert dykes' (Steinitz, 1970), and were interpreted as the

result of an increase in volume of the siliceous sediment during diagenesis. At

other horizons the ¯int is partly brecciated (Figure 12A) whereas the under- and

overlying strata are undeformed. This feature is assumed to be associated with

autobrecciation processes, also resulting from diagenetic ¯int expansion

(Kolodny, 1969; Kolodny & Garrison, 1994).

Interpretation. The hemipelagic character of this facies is indicated by the presence

of varying proportions of calcareous nannofossils and planktonic foraminifera.

The foraminiferal planktonic/benthonic ratio, which under certain conditions may

serve as a palaeobathymetric proxy (Phleger, 1964; Reiss et al., 1974; Gibson,

1989; van der Zwaan et al., 1990), has been determined routinely for all

foraminiferal samples. However, strong palaeoproductivity changes in the

southern Tethyan surface waters during this period (Almogi-Labin et al., 1993;

Eshet & Almogi-Labin, 1996) seem to exclude high resolution sea-level

reconstructions based on foraminiferal distribution patterns in this setting.

The formation of ¯int is interpreted to have been associated with marine high

productivity conditions leading to blooming of diatoms and radiolaria

(Moshkovitz et al., 1983; Reiss, 1988; Almogi-Labin et al., 1993; W. N. Krebs/

Amoco, pers. comm., 1996). Intercalation of the ¯ints in hemipelagic marls and

the presence of radiolaria suggest a hemipelagic character for most of the ¯ints.

Nevertheless, there are some exceptions, such as episodically intercalated ¯int

beds in a Coniacian mixed siliciclastic-carbonate intertidal succession (Taba,

section M, Figure 7).

6. Sequence stratigraphy

Sequence identi®cation from outcrop-scale data requires the completion of

several steps during the course of data analysis (see e.g., Van Wagoner et al.,1988). In our study, the steps were as follows: (1) Facies interpretation for all

sections was completed. (2) Horizons with facies changes were identi®ed and

vertical stacking patterns interpreted. (3) The cycles were attributed to a `high'

and a `low' order, based on their duration, amplitudes and stacking patterns. (4)

Cycles from different sections were correlated within a biostratigraphic and partly

lithostratigraphic framework.

Sequence stratigraphy in Egypt 167

Within the Turonian to Maastrichtian interval, six major sequence boundaries

of `low order' have been reconstructed (Figure 15). Differentiation between type

1 and type 2 sequences (Van Wagoner et al., 1988) is complicated in this study

Figure 7. Taba section (M), lower part. Legend in Figure 8C; for location, see Figure 1.

168 S. LuÈning et al.

because the investigated area of central Sinai represents only a small part of the

continent-basin pro®le; as a result, seismic-scale reconstruction of the whole

pro®le cannot be achieved. Low order cycle duration varies between 4 and 9 my,

which indicates a cycle order intermediate between 3rd and 2nd, according to the

schemes given in Weber et al. (1995) and Emery & Myers (1996). Maximum

error ranges for the dating of our systems tracts boundaries are estimated as � 1±

2 my. The amplitudes shown in the relative sea-level curve (Figure 15) are based

on the facies zones and unconformities as described below for the different

systems tracts in central Sinai. Facies distribution maps have been compiled for

the study area on a systems tract-scale (Figure 16A-M) based on our data and

descriptions from the literature. Higher (4th-5th) order cycles are discernible in

central-east Sinai (see below) but were not reconstructed on a basin-wide scale

because this would need a higher resolution data-base.

Sequence boundaries are numbered following a system used by Hardenbol etal. (in press). The ®rst two letters refer to the stage (e.g., Tu for Turonian), while

the letters `Sin' indicate that the sequence boundary was reconstructed for the

Sinai Peninsula. In the text we refer to the different systems tracts and their

boundaries by adding the name of the previous underlying sequence boundary,

preceded by the pre®x `post-' (for example: `post-CoSin TST' = TST between

sequence boundaries CoSin and Sa/CaSin). Further abbreviations of the

sequence stratigraphic nomenclature (LST, TST, LST, ts, sb, mfs) are explained

in Figure 8C. The sequence descriptions below are listed from the oldest to the

youngest sequence boundary.

6.1. TuSin (Turonian sequence boundary)LST (Figure 16B). In the Sheikh Attiya section (section B, Figure 6), sequence

boundary TuSin is characterized by a marked upward facies change from low

energy outer-inner shelf marls and carbonates to a thick development of silty

mud¯at facies with local sabkha gypsum formation. This signi®cant facies break is

preceded by two short-term shallowing events indicated by intercalation of

siltstone horizons into the low energy calcareous shelf facies. The stacking pattern

re¯ects prograding parasequences of the late pre-TuSin HST. Within this HST

succession (facies map in Figure 16A), an outer shelf marl was found to contain a

few planktonics, besides the dominating benthonic foraminifera, indicating the

Whiteinella archaeocretacea or Helvetoglobotruncana helvetica Zones (Figure 6). This

dating is based on the presence of Heterohelix moremani, W. archaeocretacea and

W. baltica. Calcareous nannofossils in this horizon include Lucianorhabdusquadri®dus and Quadrum gartneri, while Eiffellithus eximius is absent, suggesting an

early to middle Turonian age (L. quadri®dus Zone sensu Crux, 1982; Q. gartneriZone sensu Mortimer, 1987). The same interval yields the ammonite Choffaticerassp., which indicates a maximum age of late early Turonian (Lewy & Raab, 1976).

The sequence boundary must be only slightly younger and is estimated to be of

late early Turonian age (H. helvetica Zone).

A mid-Turonian high sea level within the H. helvetica Zone was also described

from Israel by Honigstein et al. (1989) and may be correlated with the TST/HST

below sequence boundary TuSin. From central-east Sinai, Orabi (1992, 1993)

presented evidence of the hemipelagic character of the lower part of the Lower

Turonian Abu Qada Formation which we again interpret as the pre-TuSin TST/

HST. The upper part of this formation seems, in general, to be devoid of open

marine fossils (e.g., Orabi, 1992, 1993), indicating a change of environmental

Sequence stratigraphy in Egypt 169

Figure 8. Taba section (M). 8A, upper part. B, enlargement of the interval 175-179 m from A. C,legend for Figures 5±9.

170 S. LuÈning et al.

conditions, most probably associated with the sea-level fall that resulted in the

TuSin sequence boundary. Four Cenomanian-Turonian sections and their

foraminiferal contents were described from central-west Sinai by Cherif et al.

(1989a). They documented hiatuses between the Cenomanian and the siliciclastic

Lower Turonian for Gebel Mukattab and Gebel Qabeliat, which may be related

to erosion at the TuSin sequence boundary. Although we cannot agree with some

of their sea-level interpretations, their lithological and faunal descriptions provide

valuable data to illustrate the complex lateral facies relationships within the

Turonian of Sinai, and in general support our sequence stratigraphic

interpretation. To the east, a late early Turonian omission surface was reported

from the Maktesh Ramon anticline of the Negev Foldbelt by Lewy & Avni

(1988). The authors assumed that the origin of this omission surface is associated

with "local weak differential movements" which are "superimposed on the

regional shallowing with local exposure observed in central Israel" (Lewy & Avni,

1988, p. 112). The corresponding ages suggest that the regional shallowing

mentioned may also be associated with the regional sea-level fall at the TuSin

sequence boundary.

The post-TuSin lowstand succession in the study area is characterized by

peritidal and supratidal silty mud¯at deposits in which deltaic or tidal ®ne

sandstone channels are intercalated. Locally, gypsum formed in sabkha settings.

Comparable nearshore to continental deposits are developed in neighbouring

outcrops in the southern Negev Desert. A subtropical freshwater lake facies with

an abundant angiosperm ¯ora has recently been described by Dobruskina (1997)

from the upper member of the lower Turonian Ora Shales from that region.

Abdel Gawad et al. (1992) subdivided the Turonian at Gebel Nezzazat in

central-west Sinai into three lithofacies units. While we interpret their basal

`carbonate part' as TST or HST, the middle `clastic part' may be underlain by

the TuSin sequence boundary and consequently represent the LST. Their

uppermost `carbonate part' re¯ects post-TuSin TST and HST deposits and is

most probably separated from the middle `clastic part' by a transgressive surface.

The Turonian strata in their section are overlain by Coniacian sandstones (Abdel

Gawad et al., 1992) with a basal contact interpreted here to represent the CoSin

sequence boundary. In a study from the southern Negev/central-east Sinai area,

Bartov et al. (1972; also included in Bartov & Steinitz, 1977) subdivided the

lower Turonian Ora Shales into three members. The lowest member is reported

to consist of shales and limestones with ammonites, and is interpreted here as

pre-TuSin TST. According to Bartov et al. (1972), the middle member is mainly

composed of dolomites, oolites, chert, marl, shale and sandy limestones and, in

our opinion, represents the pre-TuSin HST. The upper member varies in

composition from shaly marls to gypsum and sandstones. We interpret this unit

as the post-TuSin LST. Bartov et al. (1972) assumed that the ¯ooding of the area

with re-establishment of open marine conditions occurred during the late

Turonian with deposition of the Gero®t Formation (post-TuSin TST + HST).

Bartov et al. (1972) interpreted a shallowing at the beginning of the Coniacian

Zihor Formation which is composed of mixed carbonate-siliciclastics (post-CoSin

LST).

TST (Figure 16C). The transgressive surface (ts) in the Sheikh Attiya section (B)

is characterized by a facies change from sabkha sedimentation with gypsum-shale

interbeds to inner shelf-type green shale with oysters. The transgressive surface is

Sequence stratigraphy in Egypt 171

estimated to be of middle Turonian age (H. helvetica or M. sigali Zones) based on

the stratigraphic context. Planktonic foraminifera have not been found around

this horizon. However, the TST in sections B and M yielded ostracods of the

species Neocyprideis vandenboldi (Figures 6, 7) which represents a marker for the

Turonian (Gerry & Rosenfeld, 1973). The same horizon in section M yields rare

Spinoleberis yotvataensis which, according to Rosenfeld & Raab (1974), is also

characteristic of the Turonian.

The post-TuSin TST is developed in three of the studied sections (B, M, N1)

and is characterized by predominantly inner shelf carbonates of the low and

higher energy type. The TST interval is mainly represented by the sediments of

the Wata Formation (Gero®t Formation in Israel). The presence of higher order

sequences is marked by repeated short-term switches to peritidal conditions

indicated, for example, by silty marls with characean algae, cross-bedded ®ne

sandstones and calcareous laminites with birds-eyes. The cm/dm-scale shallowing

upward cycles in section N1 (Figure 10B) also represent higher order sequences.

In the upper part of the TST, higher order sequences are indicated by repeated

occurrences of planktonic foraminifera in the succession (e.g., upper part of TST

in N1 and B; Figure 6). The top of the TST is marked by about 10 m of

hemipelagic chalky marl (sections B, M) which re¯ect the overall deepening trend

during this transgressive phase.

HST (Figure 16D). The maximum ¯ooding surface (mfs) in the Taba section

(M) is drawn between the late TST hemipelagic chalky marls (see above) and

overlying harder calcareous marls which are probably of inner neritic origin

(Figures 7, 13B). Further to the south in the Sheikh Attiya section (B), the

maximum ¯ooding surface is represented by the boundary between a chalky

interval containing some planktonic foraminifera, and overlying, predominantly

low energy, inner shelf carbonates (Figure 6). In section M, the late TST deposits

fall within the D. primitiva Zone (the calcareous nannofossil E. eximius Zone)

which indicates an early Coniacian age (Figure 7). A few specimens of Dicarinella

sp. were also found in the ?highstand deposits of section H but the exact biozone

remains unclear. The HST of sections B (Figure 6) and H yield the ostracod

Spinoleberis yotvataensis macra, which has previously been described from the late

Turonian-Coniacian (Honigstein & Rosenfeld, 1985). In section B, the species

Ovocytheridea apiformis has been found which, according to Reyment (1960), is

characteristic of the Coniacian-Santonian (Figure 6). Within the uppermost HST

of section H, abundant occurrences of Neocyprideis vandenboldi were recorded,

which so far have only been described from the Turonian (Gerry & Rosenfeld,

1973). However, this species seems to be also present in the Lower Coniacian if

our correlations are correct. The slight discrepancy may be also related to the

problems of de®ning the Turonian-Coniacian boundary on the basis of different

fossil groups, such as ammonites and foraminifera. In conclusion, the

biostratigraphic information provided by ostracods supports the concept of an

early Coniacian age for the post-TuSin HST deposits, which in the study area are

characterized by carbonates and partly by marls of the low and higher energy

inner shelf facies. Although the HST in section M is relatively thin, a greater

thickness is present in section B.

172 S. LuÈning et al.

6.2. CoSin (Coniacian sequence boundary)LST (Figure 16E). In section M (Figure 7) the sequence boundary CoSin is

characterized by a facies change from partly phosphoritic inner shelf carbonates to

a peritidal siliciclastic system. Further to the north and west (section H; Figure 1)

the boundary is interpreted to be located where chalky HST deposits with a few

planktonic foraminifera are overlain by cross-bedded inner shelf carbonates

characterized by shallow water indicators such as ooids, ostracods and miliolid

foraminifera. The sequence boundary CoSin is of early Coniacian age (the

planktonic foraminiferal D. primitiva or D. concavata Zones and E. eximius or M.

staurophora calcareous nannofossil Zones) based on the datings of the late pre-

CoSin TST (see above) and of a higher order transgressive phase within the post-

CoSin LST (D. concavata/L. maleformis Zones, section D; Figure 9B). A roughly

coeval sea-level fall has also been described by Orabi (1991) from western Sinai.

The LST in the Taba section (section M, Figure 7) is characterized by a

subtidal to supratidal siliciclastic shallowing-upward succession (see description

above) represented by sandstone bars and peritidal deltaic silty mud¯ats, with

subordinate sabkha-bedded gypsum. Higher order sequences are indicated by

repeated intercalations of shallow marine carbonates (oyster packstones, oyster

calcareous marls in section M, and inner-outer shelf chalks in section D;

Figure 9B) into a peritidal silty mud¯at and sandstone bar facies. Relatively high

amounts of planktonic foraminifera in the chalks of section D are interpreted to

re¯ect a retrogradational stacking pattern during the latest LST. Downslope,

towards the northern part of the study area, the siliciclastics are gradually

replaced by inner shelf carbonates (e.g., in sections H, G, N2/N3). The lowstand

deposits of section H are composed of inner shelf carbonates containing the

ostracod Bythocypris windhami. This species is known from the Coniacian to

Campanian (Butler & Jones, 1957), providing further evidence for a Coniacian

age for this LST. In the middle of section G, a green algal ¯ora including

Acicularia magnapora (Figure 11E, previously described from the middle

Turonian of Abu Roash; Kuss, 1994), Halimeda sp. and Neomeris cretacea was

found. A description of higher energy oyster bioherms within a low energy inner

shelf carbonate facies, has been given by Bartov & Steinitz (1982) from the Negev

Desert and Sinai, including the Themed and Sheikh Attiya areas. A marginal

example of this facies is probably developed in section N2/N3. According to

Bartov & Steinitz (1982), the oyster bioherms developed in association with

submarine morphotectonic palaeohighs and locally inter®nger with sandstones.

In this study, higher order sequences within the post CoSin LST are developed

in section G, where siliciclastically in¯uenced carbonates and beds containing a

few planktonic foraminifera are intercalated in a calcareous inner shelf facies

succession marking short-term periods of shallowing and deepening, respectively.

This observation is consistent with results from Khalifa & Eid (1995) who carried

out a detailed facies analysis of the Coniacian-Santonian lowstand deposits of

central-east Sinai and subdivided the succession into ®ve shallowing-upward

cycles (parasequences). A typical cycle is reported to consist of (from base to top)

subtidal green glauconitic shales, carbonate packstones, phosphatic dolostones

and intertidal quartz-arenites. However, correlation of our parasequence pattern

with the cycles of Khalifa & Eid (1995) is complicated because these authors

based their model on a lithostratigraphic rather than a biostratigraphic

framework.

Sequence stratigraphy in Egypt 173

Figure 9. A, Section T2 (between Taba and Nuweiba). B, Section D (N0 Sheikh Attiya). Legend inFigure 8C; for location, see Figure 1.

174 S. LuÈning et al.

A detailed description of the late Turonian to early Santonian lithological and

facies development for southern Israel and Sinai, including the interval from the

pre-CoSin HST to post-CoSin TST, was given by Lewy (1975). His

interpretations are based on a high resolution biostratigraphic framework

provided by ammonites. Our ®ndings correspond well with many of his results

and interpretations. Lewy (1975) described a shallowing around the Turonian/

Coniacian boundary with a maximum regression during early late Coniacian

times which we correlate with sequence boundary CoSin. Renewed ¯ooding of

the shelf was inferred by Lewy (1975), based on the deposition of uppermost

Coniacian chalks. This interpretation corresponds well with the timing of the

post-CoSin transgressive surface which is discussed below. However, Lewy

mentions several tectonic events, including regional tilting and folding, which may

apply mainly for the Syrian Arc region but, according to our observations, played

no signi®cant role on the stable shelf of southern and central Sinai. Upper

Turonian and Coniacian strata are reported to be absent (Kassab & Ismael, 1994)

in the Abu Zeneima area (western Sinai). This stratal gap may be owing to

erosion and/or non-deposition at this locality during the Coniacian LST.

The north-south trend of the Coniacian palaeoshoreline in central-west Sinai,

as shown in Figure 16E, is derived from isopach and sand/shale ratio maps of the

Lower Senonian (Refaat, 1993), and is also evident in the Coniacian

palaeogeographic maps of Said (1990, p. 446). A late Turonian-early Coniacian

phase of siliciclastic facies progradation has been reconstructed for southern

Egypt (`facies 3' in Van Houten et al., 1984) and may represent the succession

from the post-TuSin LST to the post-CoSin LST. Investigations of the

geochemistry, microfacies and facies of the mixed siliciclastic-carbonate LST

sediments of the Matulla Formation were carried out by Refaat (1993) in the

Gulf of Suez region and Orabi & Ramadan (1995) in central-west Sinai.

TST and HST (Figure 16F). The TST and HST are composed of cherts and

chalks, the latter usually containing signi®cant amounts of planktonic

foraminifera. Upper parts of the sequence (Santonian) are commonly

phosphoritic. In general terms the interval corresponds to the middle and upper

non-siliciclastic part of the Sayyarim Formation in Israel. Similar age-equivalent

hemipelagites containing a rich planktonic foraminiferal fauna have been

described by Ismail (1993) from central-west Sinai. This interval has been

described from Israel by Schneidermann (1970) and others.

The post-CoSin transgressive surface, interpreted in sections M and T2

(Figures 7, 9A, 10A), separates peritidal siliciclastics and inner shelf carbonates

from overlying, probable outer shelf, cherty chalks and cherts. The age of the

transgressive surface is estimated to be the D. concavata (foraminiferal) Zone and

the L. maleformis or R. anthophorus (calcareous nannofossil) Zones. This is based

on datings of a higher order transgressive phase within the latest post-CoSin LST

(D. concavata/L. maleformis Zones in section D; see above) and of the middle

post-CoSin TST (D. concavata/R. anthophorus Zones in section M). In the Nakhl-

1 well Santonian hemipelagites of the lower D. asymetrica Zone are recorded

above sandy limestones, coinciding with a sharp drop in the resistivity curve. The

age matches the data derived from our outcrop sections. The occurrence of the

Santonian (Honigstein, 1984) ostracod Cythereis rosenfeldi rosenfeldi in the TST/

HST of section M conforms well with the dates provided by planktonic

foraminifera and calcareous nannofossils (Figure 8).

Sequence stratigraphy in Egypt 175

The exact position of the maximum ¯ooding surface is unclear. Either the HST

is lithologically similar to the TST or was eroded during the subsequent sea-level

fall. From sections M (Figure 8) and A6, the inferred maximum ¯ooding surface

is estimated to have a biostratigraphic age of D. asymetrica/L. cayeuxii Zones.

6.3. Sa/CaSin (sequence boundary near the Santonian-Campanian boundary)LST. In the study area, the sequence boundary Sa/CaSin has been found to be

well developed in three sections. In a section south of Gebel Areif El Naqa (A6)

in NE Sinai (Figure 1), the boundary is clearly marked by a pebbly intra-chalk

unconformity with a probable hiatus. The section lies on the unstable shelf.

Because of its position in a synclinal area, deposition was less in¯uenced by the

Syrian Arc compression than in the anticlinal areas. In section M (Figure 8), a

hiatus even greater than that in section A6 is developed. Sequence boundaries Sa/

CaSin and CaSin (see above) are amalgamated here, as demonstrated by a large

biostratigraphic hiatus recorded by calcareous nannofossils. In section T2

(Figure 9A), the sequence boundary is interpreted to lie between a partly silici®ed

chalk and an overlying hard calcareous marl (with oysters). The calcareous marl

contains phosphatic pebbles and ¯int fragments at the base and may represent the

lowstand deposits. In sections A6, M and T2 the sequence boundary is dated as

lying close to the Santonian-Campanian boundary (D. asymetrica-G. elevataforaminiferal zonal boundary/L. cayeuxii-B. parca nannofossil zonal boundary). A

coeval sea-level fall, marked by a clay bed intercalated in chalk throughout Israel,

has been described by Lewy (1990). Based on the absence of a shallow-water

facies, it is supposed that the post-Sa/CaSin LST was either not deposited or has

been eroded. From the Eastern Desert, Kuss & Malchus (1989) reported the

absence of Santonian-lower Campanian strata, which may be related to this sea-

level fall.

TST + HST (Figure 16G). The transgressive surface in section T2 (Figure 9A) is

interpreted to be at the lithological boundary between a hard calcareous marl with

oysters (see above) and an overlying softer chalky calcareous marl with

foraminiferal assemblages containing between 1±40% planktonic foraminifera.

Because the LST is missing in section A6, the transgressive surface here coincides

with the sequence boundary. The hemipelagic sediments just above the Sa/CaSin

unconformity in section A6 are, therefore, interpreted as representing the lower

TST. In both sections the transgressive surface is referable to the B. parca Zone

[G. elevata (foraminiferal) Zone]. The TST and HST are composed of chalk.

The position of the maximum ¯ooding surface remains unclear in the studied

sections and is estimated to be within the G. ventricosa Zone.

6.4. CaSin (Campanian sequence boundary)LST (Figure 16H). In section M, as described above, the sequence boundary

CaSin seems to be amalgamated with the Ca/SaSin sequence boundary. Here, the

boundary is characterized by an erosional lower contact between hard calcareous

marls and overlying chalky marls with a signi®cant biostratigraphic hiatus, as

demonstrated by the absence of two calcareous nannofossil biozones. The upper

unit is interpreted as a TST so that the LST in this section is missing. The

biostratigraphic ages of the post-CaSin transgressive systems tract (Q. tri®dumZone in section M, Figure 8) suggest that CaSin lies within this biozone or is

older. A bored and abraded omission surface in chert on palaeohighs in Israel of

176 S. LuÈning et al.

comparable age has been described by Lewy (1990, p. 629), and may be related

to the inferred sea-level fall recorded by the sequence boundary CaSin. Moreover,

freshwater diagenesis of these cherts has been reported by Kolodny et al. (1980).

TST (Figure 16I). In section M, the TST is represented by calcareous marls and

¯ints which are partly interbedded. The post-CaSin transgressive surface is within

the G. ventricosa foraminiferal and Q. tri®dum or younger nannofossil zones. The

TST may be correlated with the lower part of the Duwi Formation (e.g., in

Orabi, 1991).

HST (Figure 16K). The upper Campanian succession of section M contains

several tempestites composed of inner shelf bioclastic phosphorites intercalated

with calcareous marls. The occurrence of the tempestites is interpreted to be

related to a gradual fall in relative sea level leading to a northward progradation of

the tempestite facies. For section M, the maximum ¯ooding surface is assumed to

lie below the ®rst signi®cant occurrence of these tempestites. That reworking

plays a signi®cant role in phosphorite formation in the Campanian of the region

was previously noted by, for example, Kolodny (1980), Kolodny & Garrison

(1994) and Soudry (1987, 1992). From section M it is inferred that the

maximum ¯ooding surface lies within the Q. tri®dum Zone and probably within

the G. calcarata Zone.

The TST and HST of section M were signi®cantly in¯uenced by high

productivity conditions, as indicated by the widespread development of chert and

phosphorite. High productivity conditions prevailed over most of the southern

Tethys region during the Santonian to early Maastrichtian (e.g., Kolodny &

Garrison, 1994). Phosphatic omission surfaces from the Upper Campanian of

southern Israel have been described by Soudry & Lewy (1990). They may be

attributed to autocyclic changes in the energy regime rather than to major sea-

level falls.

6.5. Ca/MaSin (sequence boundary near the Campanian-Maastrichtian boundary)LST (Figure 16L). The sequence boundary Ca/MaSin in section M is

characterized by a drastic facies change from inner shelf bioclastic phosphorites to

probable marshy greyish-red marls devoid of microfossils, but with abundant

plant fragments (Figure 10C). The sequence boundary is assumed to lie within

the Q. tri®dum Zone and around the G. calcarata-G. falsostuarti zonal boundary.

Just below the sequence boundary within the latest HST, the ammonite

Solenoceras humei was found which is interpreted by Abdel-Gawad (1990) to

indicate the latest Campanian. Hiatuses around the Campanian-Maastrichtian

boundary are also described from Israel (Shiloni et al., 1988; Almogi-Labin et al.,1990; Lewy, 1990), from western Sinai (Orabi, 1991), and from the Eastern

Desert (Luger, 1988; Hendriks & Luger, 1987). In Israel, freshwater diagenesis

around the Campanian-Maastrichtian boundary was mentioned by Reiss (1988).

Detailed sedimentological investigations of the late Campanian-early

Maastrichtian phosphorites of Egypt (excluding Sinai) were carried out by Glenn

(1990). He proposed a sequence stratigraphic development for this period which

is similar to our model. Small differences in the stratigraphic position of the

sequence boundary around the Campanian-Maastrichtian boundary (latest

Campanian interpreted by Glenn vs earliest Maastrichtian in our study) may be

largely attributed to the uncertainties involved in dating by planktonic

foraminifera and other fossil groups (see discussion above).

Sequence stratigraphy in Egypt 177

Figure 10. Supratidal, calcareous peritidal and low energy carbonate inner shelf facies. A, Post-CoSin transgressive surface separating (hard) inner shelf-type grainstones (lower TST) from(soft) underlying laminated reddish-greenish supratidal shales, siltstones and ®ne sandstones(upper LST), hammer at central left for scale; Upper Coniacian, section M (Taba, transgressivesurface at 69 m in Figure 7). B, High frequency sequences consisting of (laminated) tidalites (td)and subtidal bioturbated horizons (st); Upper Turonian of section N1. C, Supratidal marl withabundant plant fragments close to the Campanian-Maastrichtian boundary representing the postCa/MaSin LST deposits in section M (Taba, at 177 m in Figure 8). D, Upper Turonian peritidalstromatolite in section N1. E, F, Fenestrate fabrics (E) and inverse grading (F) in peritidal peloi-

178 S. LuÈning et al.

TST and HST (Figure 16 M). The transgressive surface in section M is

interpreted to lie between the marl deposits rich in plant fragments and overlying

phosphoritic inner shelf carbonates. The transgressive surface is dated as Q.tri®dum Zone/G. falsostuarti Zone. While the lower TST is dominated by

phosphoritic inner shelf carbonates, the upper part of the TST and the HST are

composed of chalks and chalky calcareous marls containing abundant planktonic

foraminifera and calcareous nannofossils (sections M, C, F, P). In section P, a

hemipelagic sponge fauna occurs, including hexactinellids of the genera

Rhizopoterion and Schizorabdus which were previously described by Herrmann-

Degen (1980) from ?Maastrichtian-Paleocene chalks of the Egyptian Western

Desert.

A pronounced ¯ooding event around the Campanian-Maastrichtian boundary

was also inferred from sections in central-west Sinai (Cherif et al., 1989b) where

the transgressive surface separates siliciclastics of the Duwi Formation from

overlying chalks with planktonic foraminifera. Identi®cation of the maximum

¯ooding surface in this hemipelagic setting is complicated because of the

monotonous lithological record. In addition, it seems that foraminifera cannot be

used for sea-level reconstruction because of the strong in¯uence of

palaeoproductivity on their abundance distribution during the Santonian to

Maastrichtian (Almogi-Labin et al., 1993; Eshet & Almogi-Labin, 1996).

Nevertheless, a late Maastrichtian age for the maximum ¯ooding surface is

assumed on the basis of the sequence context.

6.6. Ma/DaSin (sequence boundary just above the Maastrichtian-Danian boundary)Sequence boundary. The sequence boundary Ma/DaSin in the study area was

interpreted to lie just above the K/T boundary and is marked in sections C and P

by the presence of reworked Cretaceous planktonic foraminifera in lower

Paleocene (NP1) marls with high foraminiferal planktonic values. The Paleocene

marls are interpreted to represent the post-Ma/DaSin TST and HST with typical

shallow marine LST deposits missing (LuÈning et al., 1998). In section M, the late

Maastrichtian Abathomphalus mayaroensis Zone is absent, which may be

interpreted to re¯ect local deep erosion during the Ma/DaSin sea-level fall.

Further evidence of sea-level fall just above the K/T boundary comes from a

number of other studies in the region. For example, very shallow conditions are

reported from the K/T interval in central-west Sinai by Ismail (1992) and Abbass

et al. (1994) and in the Negev by Keller & Benjamini (1991) and Eshet et al.(1992). A hiatus at the K/T boundary has been described from western Sinai

from the Abu Zeneima area (Anan 1992; Marzouk & Hussein, 1994), Gebel

Mokattab (Ayyad & Hamama, 1991), Wadi Feiran, Gebel Qabeliat (Marzouk &

Abou-El-Enein, 1995) and several subsurface sections (Orabi, 1991). Reworked

Cretaceous foraminifera have been observed within the Parvularuglobigerinaeugubina/Parasubbotina pseudobulloides Zones in central Egypt (Luger et al., 1997,

dal carbonates from the Upper Turonian of section B (at 184 m in Figure 6); photograph fromthin section, sample B1-18. G, Dolomitic wackestone with Acicularia (centre), serpulids (arrow)and molluscan shells; photograph from thin section, section G, sample G1-7. H, Dolomiticmicrite with oyster shells; photograph from thin section, section G, sample G1-7. I, Dolomiticmiliolid (upper arrow) wacke/packstone with reworked micritic clasts (lower arrow), transitionbetween low and high energy calcareous inner shelf facies, photograph from thin section, sectionG, sample G1-5.

Sequence stratigraphy in Egypt 179

Figure 11. High and low energy calcareous inner shelf facies. A, Thalassinoides traces (horizonabove hammer) in the Turonian low energy carbonate inner shelf facies of section B (at 128 m inFigure 6). B, Cross-strati®ed packstones with basal scours (to right of ®nger-tip) cut into under-lying calcareous marl (Upper Turonian of section B, at 122 m in Figure 6). C, Wave ripples inoolitic packstone (Upper Turonian of section H). D, Rip-up clast in cross-strati®ed dolomiticbioclastic packstone (Upper Turonian of section B, at 136 m in Figure 6). E, Well sorted dolo-mitic packstone with Acicularia magnapora (arrow) (section G, photograph from thin section G1-8). F, High energy oolite with intercalated mud layers and relict mud drapes in ripple troughswhich were deposited during low energy periods (Upper Turonian of section B, at 161 m inFigure 6, negative print from sample B1-15). G, Oolitic grainstone with large reworked micriticclasts (Upper Turonian of section B, at 163 m in Figure 6; photograph from thin section B1-16).H, Grainstone with abundant micritic clasts and Dicyclina sp. (®lament structures) (Upper Turo-nian of section H, photograph from thin section H1-5).

180 S. LuÈning et al.

Figure 12. Chert and siliciclastic peritidal facies. A, Autobrecciation in Campanian chert caused bydiagenetic expansion (section M, at 160 m in Figure 8). B-D, Slumped horizon with slump folds(B, C) and slump shear planes (C) intercalated in ®nely laminated undeformed ®ne sandstones(D); Coniacian, post-CoSin LST in section M (at 52 m in Figure 7). E, Cross-bedded mediumsandstones in the Coniacian of section M (at 55 m in Figure 7). F, Coarse sandstone pseudo-nodules which have sunk into ®ne sandstones as a result of loading (see Allen, 1982, p. 360);Coniacian of section M (at 94 m in Figure 7). G, Phosphoritic quartzose coarse sandstone, quartzgrains subangular, dark grains phosphoritic; Coniacian-Santonian in section T2 (at 5 m inFigure 9A); photograph from thin section T2-3. H, Thin alternations of dolomitic micrite andrippled ®ne- to medium-grained quartz sandstone; Coniacian of section M (at 76 m in Figure 7);photograph from thin section M2-25. I, Micrite with abundant angular quartz grains and mol-luscan shells; Coniacian of section M (at 71 m in Figure 7); photograph from thin section M2-23.

Sequence stratigraphy in Egypt 181

in press). The subsequent Paleocene sea-level development for central-east Sinai

has recently been described by LuÈning et al. (1998).

7. Comparison with sea-level reconstructions from neighbouring regions

7.1. Egypt, Israel and JordanThe relative sea-level curve proposed here for central-east Sinai (Figure 15)

corresponds well with those constructed for Sinai and Israel by Reiss (1988) and

Kuss & Bachmann (1996) (Figure 17). A good correlation also exists with the

cycle chart for Israel given by Flexer et al. (1986) except for the sea-level fall

associated with sequence boundary CaSin, which is missing in their model

(Figure 17). Several signi®cant mismatches were observed in the correlation of

our reconstructions with the `Egyptian eustatic sea-level' chart by Said (1990). A

more complex sea-level history compared to that observed in central-east Sinai

was described by Lewy (1990), mainly from the unstable shelf of Israel

(Figure 17). Differences may be explained by tectonic movements of the Syrian

Arc. However, the main cycles correlate well with our sea-level model for central-

east Sinai. A sequence stratigraphic model for the Upper Cretaceous to Eocene of

Jordan has recently been elaborated by Powell et al. (1996). Because their

approach does not include biostratigraphic ages for their systems tracts, a direct

correlation with our model is dif®cult. Nevertheless, their regional sea-level

reconstruction also seems to correspond broadly to the situation encountered in

central Sinai.

7.2. Arabian PeninsulaThe Cretaceous sea-level charts for the Arabian Peninsula (Harris et al., 1984)

and the United Arab Emirates (Alsharhan & Kendall, 1995) in general differ from

the sea-level history in central Sinai because of strong tectonic activity in Arabia

Figure 13. Exposures of section at Taba. A, Coniacian to Campanian succession. Siliciclastics domi-nate the Coniacian, whereas a more hemipelagic facies with cherts and (partly silici®ed) chalksprevails from the Santonian to Campanian; corresponds to 130-160 m in Figure 8A. B, Signi®-cant ¯ooding surface (fs) within the post TuSin-TST between Upper Turonian oolites and softmarly outer shelf deposits containing planktonic foraminifera. The sequence boundary lies inthe uppermost part of the illustrated succession and is overlain by siliciclastics; section M (fs at32 m in Figure 7).

182 S. LuÈning et al.

during that time. An exception is the K/T boundary for which these authors

described the development of a pronounced unconformity.

7.3. TunisiaA sequence stratigraphic study of the Turonian of west-central Tunisia has

recently been presented by SaõÈdi et al. (1997). They subdivided the Turonian into

two sequences, one of late Cenomanian-middle Turonian, the other of middle -

late Turonian age. The sequence boundary separating the two sequences is

reported to lie within the Helvetotruncana helvetica Zone, and the upper boundary

of the second sequence is characterized by an early Coniacian shallowing event

(SaõÈdi et al., 1997). Timing of these sequences corresponds well with the

reconstructions for central-east Sinai and yields evidence of inter-regional

processes controlling Turonian sedimentation in the southern Tethys. However,

the facies in central-west Tunisia seems to be slightly deeper than in central-east

Sinai because the LST is described as an inner shelf deposit and the TST/HST as

containing signi®cant pelagic elements (SaõÈdi et al., 1997), contrasting with the

Figure 14. Aa-c, Ba-b, Choffaticeras sp.; lower Turonian in section B (at 13 m in Figure 6). C, Sole-noceras humei, Campanian-Maastrichtian boundary just below the sequence boundary Ca/MaSin(section M, at 176.5 m in Figure 8A, B). D, E, Porifera from the late Maastrichtian Abathompha-lus mayaroensis Zone in section P; specimen in D in situ, specimens in E moved for photograph.

Sequence stratigraphy in Egypt 183

peri- to supratidal post-TuSin LST and mainly inner shelf post-TuSin TST/HST

in Sinai.

7.4. Eustatic sea-level curveIt is standard procedure in many regional studies to compare the interpreted

regional sea-level cycles with the eustatic model published by Haq et al. (1987).

However, today many stratigraphic intervals covered by the Exxon-curve,

especially for the Late Cretaceous, are questioned to some degree (see Hancock,

1989, 1993; Miall, 1991, 1992). Because of the apparent restrictions of the

concept of a single global eustatic curve it may be better in most cases to consider

several regional curves. A comparison of the major sea-level ¯uctuations of Sinai

with those described in Haq et al. (1987) shows that the cyles in Sinai are in most

cases longer than the 3rd order cycles in the `eustatic' sea-level chart. Considering

Figure 15. Regional sequence chart and relative sea-level curve for central-east Sinai. No numericalscale for the sea-level curve is implied. Sequence boundary names consist of an abbreviation ofthe stage followed by the suf®x `Sin'. Stage boundary ages are based on Gradstein et al. (1995),correlation with planktonic foraminiferal biozones on Bralower et al. (1995), and correlation offoraminiferal and nannofossil zones on our data and those in Bralower et al. (1995) and Eshet &Moskovitz (1995) (see text for discussion); ts � transgressive surface, mfs �maximum ¯oodingsurface.

184 S. LuÈning et al.

the unavoidable uncertainties in sequence timing, for each sequence in Sinai

several potential `eustatic' sequences can be found; hence, convincing correlations

can seldom be accomplished. This applies especially to the sequence boundaries

in the Lower Turonian (TuSin), around the Santonian-Campanian boundary (Sa/

CaSin), and in the middle Campanian (CaSin) (Figure 17). Higher probabilities

for synchroneity exist for the sequence boundaries of the Lower Coniacian

(CoSin), at the Campanian-Maastrichtian boundary (Ca/MaSin) and just above

the K/T boundary (Ma/DaSin) (Figure 17). A certain mismatch in the timing of

the CoSin sequence boundary may be related to differences in the lower limit of

the D. primitiva Zone which in the chart of Haq et al. (1987) lies within the

Upper Turonian and in a scheme of Bralower et al. (1995), at the Turonian-

Coniacian boundary.

In a `eustatic' sea-level reconstruction based on data from the tectonically

stable Russian Platform and Siberia, Sahagian et al. (1996) described pronounced

mid Turonian and mid Coniacian `eustatic' sea-level falls which may be

correlated with sequence boundaries TuSin and CoSin, respectively, in central-

east Sinai.

8. Conclusions

The deposition of the Upper Cretaceous of central-east Sinai was controlled by a

long-term transgressive phase and several higher order sea-level ¯uctuations. In

our study, six different facies zones have been differentiated. The main

depositional elements are coastal mud¯ats with tidal channels, gypsiferous sabkha

plains, peritidal siliciclastics, peritidal carbonates, high and low energy carbonate

inner shelf facies, and microfossil-rich outer shelf pelites. The paper presents the

®rst sequence stratigraphic interpretation for this interval in the region based on

detailed sedimentological, biostratigraphical and palaeoecological investigations of

13 Turonian-Maastrichtian sections and reviews all available data from the

literature. The biostratigraphy is based mainly on planktonic foraminifera,

calcareous nannofossils, ostracods and ammonites. The study is restricted to an

area which was tectonically quiescent during the Late Cretaceous-early Tertiary,

lying to the south of the Syrian Arc intraplate foldbelt which experienced major

uplift during this period. Within the Turonian to Maastrichtian interval, six major

sequence boundaries have been recognized. Cycle duration varies between 4 and

9 Ma, which points to a cycle order intermediate between 3rd and 2nd.

Correlation with other sea-level reconstructions for the region (Egypt, Israel,

Jordan, Tunisia) points to more or less synchronous regional sea-level changes.

Comparison of the regional sequences with the `eustatic' model of Haq et al.(1987) involves uncertainties; nevertheless some of the sea-level ¯uctuations

recorded in Sinai may be attributed to worldwide eustatic changes.

Acknowledgements

We thank the following persons for various kinds of assistance during the course

of our study: Peter Luger, R. Gietl, M. Bachmann (University of Bremen), M.

Brinkmann, E. Friedel, R. Kunow, J. Beyer and K. Virkus (technical assistance),

A. Bassiouni, A. Strougo (Ain Shams University, Cairo), W. R. Fitches

(University of Wales, Aberystwyth), R. E. Garrison (University of California,

Santa Cruz), Y. Kolodny (Hebrew University, Jerusalem) and Christian

Sequence stratigraphy in Egypt 185

Figure 16. A-M, Palaeogeographic maps for central Sinai. Legend in (H). Letters indicate datapoints from this study and from the regional literature: a, this study; b, Orabi (1992, 1993); c,Cherif et al. (1989a); d, Orabi (1991); e, Abdel Gawad et al. (1992); f, Bartov et al. (1972); g,Dobruskina (1997); h, Shahin & Kora (1991); i, Ziko et al. (1993); k, Khalifa & Eid (1995); l,Enani et al. (1994); m, Refaat (1993); n, Bartov & Steinitz (1982); p, Orabi & Ramadan (1995);q, Kassab & Ismael (1994); r, Ismail (1993); s, Kora & Genedi (1995); t, Flexer (1968); u,

186 S. LuÈning et al.

Cherif et al. (1989b). Additional data concerning the strike of facies belts are from isopach mapsin Bartov & Steinitz (1977) in maps A, I and K. Palaeogeographic data from Lewy (1975) areincluded in maps C±E. The NW-SE strike of the coastline in west Sinai in map E is documen-ted, e.g., in Said (1990, p. 446), Refaat (1993) and Enani et al. (1994). The strike of facies beltsin map H is based on Said (1990, p. 448). Additional palaeogeographic data in map M comesfrom Camoin et al. (1993).

Sequence stratigraphy in Egypt 187

Figure 17. Correlation of the relative sea-level curve for central-east Sinai (this study, position ofsequence boundaries indicated) with other regional sea-level reconstructions as well as the`eustatic' model proposed by Haq et al. (1987); see text for discussion. For better comparison,the stage and biozonal scheme from Haq et al. (1987) has been retained. Another scheme whichhas been corrected for absolute ages and biozone correlation as recognized in the study area, ispresented in Figure 15. The sea-level reconstructions of Flexer et al. (1986), Reiss (1988), Lewy(1990), Said (1990), and Kuss & Bachmann (1996) in general were not developed within a highresolution biostratigraphic planktonic foraminiferal and calcareous nannofossil framework so thatcorrelation between the different curves is uncertain.

188 S. LuÈning et al.

Neumann (Free University of Berlin). Data on the Nakhl-1 well was made

available by the Gulf of Suez Petroleum Company, Cairo. Funding for S.L. and J.

K. was generously provided by the Deutsche Forschungsgemeinschaft (DFG

grant Ku 642/8-1). The reviews and comments by J. R. Ineson (Geological

Survey of Denmark and Greenland), D. J. Batten (University of Wales,

Aberystwyth) and an anonymous colleague greatly improved the manuscript.

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Appendix

List of species referred to in the text

Planktonic foraminiferaAbathomphalus mayaroensis (Bolli, 1951)Archaeoglobigerina blowi Pessagno, 1967Archaeoglobigerina bosquensis Pessagno, 1967Archaeoglobigerina cretacea (d'Orbigny, 1840)Dicarinella asymetrica (Sigal, 1952)Dicarinella canaliculata (Reuss, 1854)Dicarinella concavata (Brotzen, 1934)Dicarinella primitiva (Dalbiez, 1955)Gansserina gansseri (Bolli, 1951)Globigerinelloides ultramicra (Subbotina, 1949)Globotruncana aegyptiaca Nakkady, 1950Globotruncana angulata Nakkady, 1950Globotruncana arca (Cushman, 1926)Globotruncana bulloides Vogler, 1941Globotruncana falsostuarti Sigal, 1952Globotruncana orientalis El Naggar, 1966Globotruncana ventricosa White, 1928Globotruncanita calcarata (Cushman, 1927)Globotruncanita conica (White, 1928)Globotruncanita elevata (Brotzen, 1934)Globotruncanita pettersi (Gandol®, 1955)Globotruncanita stuartiformis (Dalbiez, 1955)Globotruncanita subspinosa (Pessagno, 1960)Hedbergella delrioensis (Carsey, 1926)Hedbergella holmdelensis Olsson, 1964Hedbergella simplex (Morrow, 1934)Heterohelix moremani (Cushman, 1938)Heterohelix reussi (Cushman, 1938)Marginotruncana marginata (Reuss, 1845)Marginotruncana pseudolinneiana Pessagno, 1967Marginotruncana renzi (Gandol®, 1942)Marginotruncana schneegansi (Sigal, 1952)Marginotruncana sigali (Reichel, 1950)Marginotruncana sinuosa Porthault, 1970Plummerita reicheli (BroÈnnimann, 1952) [sensu Masters, 1993]Rosita plummerae (Gandol®, 1955)Rugoglobigerina hexacamerata BroÈnnimann, 1952Rugoglobigerina rugosa (Plummer, 1926)

194 S. LuÈning et al.

Rugotruncana subcircumnodifer (Gandol®, 1955)Rugotruncana subpennyi (Gandol®, 1955)Whiteinella archaeocretacea Pessagno, 1967Whiteinella baltica Douglas & Rankin, 1969Whiteinella inornata (Bolli, 1957)Whiteinella paradubia (Sigal, 1952)

Calcareous nannofossilsArkhangelskiella cymbiformis Vekshina, 1959Aspidolithus parcus constrictus (Hattner et al., 1980) Perch-Nielsen, 1984Aspidolithus parcus expansus (Wise & Watkins in Wise, 1983) Perch-Nielsen, 1984Aspidolithus parcus parcus (Stradner, 1963) NoeÈl, 1969Calculites obscurus (De¯andre, 1959) Prins & Sissingh in Sissingh, 1977Ceratolithoides aculeus (Stradner, 1961) Prins & Sissingh in Sissingh, 1977Cribrosphaerella ehrenbergii (Arkhangelsky, 1912) De¯andre in Piveteau, 1952Eiffellithus eximius (Stover, 1966) Perch-Nielsen, 1968Eiffellithus turriseiffelii (De¯andre in De¯andre & Fert, 1954) Reinhardt, 1965Eprolithus ¯oralis (Stradner, 1962) Stover, 1966Glaukolithus diplogrammus (De¯andre in De¯andre & Fert, 1954) Reinhardt, 1964Kamptnerius magni®cus De¯andre, 1959Lithraphidites carniolensis De¯andre, 1963Lithraphidites praequadratus Roth, 1978Lithraphidites quadratus Bramlette & Martini, 1964Lucianorhabdus arcuatus Forchheimer, 1972Lucianorhabdus cayeuxii De¯andre, 1959Lucianorhabdus maleformis Reinhardt, 1966Lucianorhabdus quadri®dus Forchheimer, 1972Manivitella pemmatoidea (De¯andre in Manivit, 1965) Thierstein, 1971Microrhabdulus decoratus De¯andre, 1959Micula concava (Stradner in Martini & Stradner, 1960) Verbeek, 1976Micula decussata Vekshina, 1959Micula murus (Martini, 1961) Bukry, 1973Micula prinsii Perch-Nielsen, 1979Micula staurophora (Gardet, 1955) Stradner, 1963Prediscosphaera cretacea (Arkhangelsky, 1912) Gartner, 1968Prediscosphaera spinosa (Bramlette & Martini, 1964) Gartner, 1968Prediscosphaera stoveri (Perch-Nielsen, 1968) Sha®k & Stradner, 1971Quadrum gartneri Prins & Perch-Nielsen in Manivit et al., 1977Quadrum sissinghii Perch-Nielsen, 1984Quadrum tri®dum (Stradner in Stradner & Papp, 1961) Prins & Perch-Nielsen in Manivit et al., 1977Reinhardtites anthophorus (De¯andre, 1959) Perch-Nielsen, 1968Reinhardtites levis Prins & Sissingh in Sissingh, 1977Rhagodiscus angustus (Stradner, 1963) Reinhardt, 1971Stradneria crenulata (Bramlette & Martini, 1964) NoeÈl, 1970Tetrapodorhabdus decorus (De¯andre in De¯andre & Fert, 1954) Wind & Wise in Wise & Wind, 1977Tranolithus manifestus Stover, 1966Tranolithus phacelosus Stover, 1966Watznaueria barnesae (Black in Black & Barnes, 1959) Perch-Nielsen, 1968Zeugrhabdotus embergeri (NoeÈl, 1959) Perch-Nielsen, 1984Zeugrhabdotus pseudanthophorus (Bramlette & Martini, 1964) Perch-Nielsen, 1984

OstracodsBythocypris windhami Butler & Jones, 1965Cythereis gr. ovata (Roemer, 1841) Rosenfeld & Raab, 1974Cythereis rosenfeldi rosenfeldi Honigstein, 1984Neocyprideis vandenboldi Gerry & Rosenfeld, 1973Ovocytheridea apiformis Reyment, 1960Spinoleberis yotvataensis macra Honigstein & Rosenfeld, 1985Spinoleberis yotvataensis yotvataensis Rosenfeld, 1974 in Rosenfeld & Raab, 1974

Calcareous algaeAcicularia magnapora Kuss, 1994Neoneris cretacea (Deloffre, 1970)

AmmonitesSolenoceras humei (Douville, 1928)

EchinoidsHemiaster solignaci Lambert, 1932

Sequence stratigraphy in Egypt 195

OystersGyrostrea antwani Malchus, 1990Pycnodonte vesiculare (Lamarck, 1806)

List of biozones referred to in the text, in stratigraphic order from oldest to youngest

Planktonic foraminiferaParasubbotina pseudobulloides Zone (Leonov & Alimarina, 1961)Parvularuglobigerina eugubina Zone (Luterbacher & Premoli Silva, 1964)Abathomphalus mayaroensis Zone (BroÈnnimann, 1952)Gansserina gansseri Zone (BroÈnnimann, 1952)Globotruncana falsostuarti Zone (Robaszynski et al., 1984)Globotruncanita calcarata Zone (Herm, 1962)Globotruncana ventricosa Zone (Dalbiez, 1955)Globotruncanita elevata Zone (Postuma, 1971)Dicarinella asymetrica Zone (Postuma, 1971)Dicarinella concavata Zone (Sigal, 1955)Dicarinella primitiva Zone (Caron, 1978)Marginotruncana sigali Zone (Barr, 1972)Helvetoglobotruncana helvetica Zone (Sigal, 1955)Whiteinella archaeocretacea Zone (Bolli, 1966)

Calcareous nannofossilsMicula prinsii Zone (Perch-Nielsen, 1979; emend. Romein & Smit, 1981)Micula murus Zone (Martini, 1969; emend. Perch-Nielsen et al., 1982)Lithraphidites quadratus Zone (Cepek & Hay, 1969; emend. Bukry & Bramlette, 1970)Arkhangelskiella cymbiformis Zone (Perch-Nielsen, 1972; emend. Martini, 1976)Quadrum tri®dum Zone (Bukry & Bramlette, 1970)Broisonia parca Zone (Sissingh, 1977; emend. Crux, 1982)Lucianorhabdus cayeuxii Zone (Sissingh, 1977; emend. Crux, 1982)Reinhardtites anthophorus Zone (Sissingh, 1977)Lucianorhabdus maleformis Zone (Crux, 1982)Micula staurophora Zone (Sissingh, 1977; emend. Crux, 1982)Eiffellithus eximius Zone (Verbeek, 1976; emend. Crux, 1982)Quadrum gartneri Zone (Cepek & Hay, 1969; emend. Manivit et al., 1977)Lucianorhabdus quadri®dus Zone (Crux, 1982)

196 S. LuÈning et al.