Sinai: sequence stratigraphy and facies distribution

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Geological Society, London, Special Publications

The Middle Cretaceous carbonate ramp of the northern Sinai: sequence stratigraphy and facies distribution Martina Bachmann and Jochen Kuss Geological Society, London, Special Publications 1998, v.149; p253-280. doi: 10.1144/GSL.SP.1999.149.01.13

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The Middle Cretaceous carbonate ramp of the northern Sinai: sequence stratigraphy anti facies distribution MARTINA

BACHMANN

& JOCHEN

KUSS

1University o f Bremen, Fachbereich Geowissenschafien, P. O. Box 330 440, 28334 Bremen, Germany (e-maik bachmann@geochron, uni-bremen, de) Abstract: The Middle Cretaceous carbonate ramp of the northern Sinai developed in two

stages on a passive continental margin overlying basal rift suites. The lower delta-dominated ramp (Upper Aptian-Lower Albian) in the north evolved into the upper carbonatedominated ramp (Middle Albian-Cenomanian) which covered the central and southeastern parts of the Sinai. The transition between the two ramp settings took place during a period of a second-order sea-level rise and a major change in climatic conditions. The sedimentary patterns are superimposed upon by higher-frequency relative sea-level changes. The influence of the third-order relative sea-level changes on the ramp deposition was reconstructed on the basis of facies patterns, sedimentary geometries, and the distribution of microfacies types. As regional tectonic movements were of subordinate importance, a sequence stratigraphic interpretation allows a fine-scale estimation of the changing ramp settings and their characteristics. The combined use of semiquantitative microfacies analysis and sequence stratigraphy allows study of the factors controlling deposition during the different systems tracts, including the respective microfacies distributions. These factors indicate that third-order sea-level fluctuations result not only in simple shifts of facies belts up and down the ramp but also in changing environmental factors such as water circulation, carbonate production and siliciclastic input. Relative sea-level changes are one of the main factors controlling the sedimentation patterns on carbonate ramps. Because of the Cretaceous greenhouse situation, long-term third-order sealevel changes are the most prominently recorded relative sea-level fluctuations (Tucker et al. 1993). They are easily discernible within the Upper Aptian to Cenomanian carbonates of the northern Sinai, which were formed at the southern passive margin of the Tethyan Ocean during times when major tectonic movements could be excluded. Moreover, the excellent exposures of the Middle Cretaceous carbonate ramp of the northern Sinai favour studies of facies development and sequence stratigraphy. Their interplay was controlled by various factors, among which we assume relative sea-level changes to have been of major importance. These changes were reconstructed from sedimentological analysis. In contrast to many ramp models, where sealevel changes are expressed by simple shifts of the facies belts up and down the ramp (e.g. Tucker et al. 1993; Handford & Loucks 1993), the detailed microfacies analysis used here allows us to demonstrate more complex reactions of the various ramp environments to sealevel changes. The bio- and lithostratigraphic framework for the shallow shelf area is based on 15 stratigraphic sections. The ramp models documented here were reconstructed by means of the most indicative nine sections, reflecting a S - N

transect from continental to shallow marine environments, whereas d e e p e r marine sediments are not present in outcrops (Fig. 1). This transect spans a 55 km wide segment of the Middle Cretaceous shelf, running perpendicular to its shoreline. A detailed bed-by-bed analysis enabled us to reconstruct lateral and vertical shifts of sedimentation patterns with respect to gradual changes of the ramp palaeoenvironments. Two ramp systems can be discerned: an older siliciclastically influenced ramp (Upper Aptian to Lower Albian, unit A), and a younger mainly carbonate ramp (Middle Albian to Cenomanian, unit B). In particular, the correlations between the six sections of the central and northern parts (sections 1-6, Fig. 1) provide excellent evidence for these patterns. Most of them were also determined in the southern exposed sections (7-9, Fig. 1). Here, the stratigraphic record of the marine units begins later than in the north owing to a second-order sealevel rise (Kuss & B a c h m a n n 1996), which means that only the uppermost part of unit A is represented in two sections of the southern part (see Fig. 5, below). Here, frequent intercalations of continental siliciclastic deposits within the later c a r b o n a t e - d o m i n a t e d r a m p (unit B) demonstrate the retrogradation of the nearshore setting. The transition between the two ramp systems during Early to Mid-Albian times is evident in both areas, which have been correlated by means of biostratigraphy and sequence

BACHMANN,M. & KUSS,J. 1998. The Middle Cretaceous carbonate ramp of the northern Sinai: sequence stratigraphy and facies. In: WRIGHT,V. R & BURCHETrE,T. E (eds) CarbonateRamps. Geological Society, London, Special Publications, 149, 253-280.


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M. BACHMANN & J. KUSS

stratigraphy. These results are used to demonstrate the changing palaeogeography during the different systems tracts and to establish different depositional models for each systems tract, allowing us to consider palaeoenvironmental changes with respect to the third-order sea-level changes.

Geological framework The marine Cretaceous strata of the Sinai were formed along the northern rim of the passive continental margin of northeast Africa. Here, sedimentation took place on a widely extended shelf system and was influenced by major tectonic events during Mesozoic-Early Tertiary times, coeval with the opening and closure of the Neotethys Ocean (Keeley 1994). Extensional rift-induced activities during Jurassic to Early Cretaceous times were due to an E - W propagating rift system north of the recent Sinai

coastline, resulting in a progressively northward subsiding passive margin, along NE-SW trending, NW dipping faults cross-cutting the northern and central parts of the Sinai (Keeley 1994, Ayyad & Darwish 1996). During Late Cretaceous-Early Tertiary times, these faults were rejuvenated under a compressional stress field and formed the 'unstable shelf' (Krenkel 1924). As a result, the NE-SW trending foldbelts of the Syrian Arc evolved, extending from Syria and crossing the unstable shelf of the central Sinai (Shahar 1994) to the Gulf of Suez area (Kuss & Bachmann 1996). The Late Jurassic-Early Cretaceous tectonic movements along the Levantine coast induced large fault-controlled half-grabens within the unstable areas. However, the imprints of halfgraben structures could not be demonstrated from the sedimentary record studied; therefore it is assumed that the Middle Cretaceous sections considered here were formed during

30"

J

5 Location of the Sections Jurassic ' l

' Aptian and Albian Cenomanian to Turonian

,~ ~) ~' @ ~

Rizan Aneiza (RN) G. El Mistan North (Mi) G. El Mistan South (MiS) G. Raghawi (R) G. Mansoura North (M)

~ ~ (~ (9) "

G Mansoura East (ME) G. Halat (H) G. Minshera (GM) G. Areif el Naqa (AN1,2) " Traverse Fig. 9

Fig. 1. Simplified geological map of the northern Sinai, showing the location of the Aptian to Cenomanian sections mentioned in the text. The sections include a transect along the inner ramp.


MIDDLE CRETACEOUS SINAI RAMP periods of relative tectonic quiescence that favoured the establishment of an extended carbonate shelf system, which was mainly influenced by relative sea-level changes and terrestrial input. The stepwise encroachment of the Late Aptian-Turonian sea towards the south led to the marine cover of the former continental areas (Fig. 2) and in a consequently northward thickening sedimentary succession (Kuss & Bachmann 1996). Based on new geophysical interpretations of the northward offshore sedimentary cover, Hirsch et al. (1995) argued that a continuation of the continental crust, including a thick sedimentary cover of Phanerozoic sediments, has to be expected further offshore. This implies a p r o l o n g a t i o n of the Cretaceous onshore carbonate shelf deposits further to the north and a more northward position of the shelf hinge line within the area of today's Eastern

255

Mediterranean Basin (Cohen et al. 1990). A second consequence concerns the extension and morphology of the Middle Cretaceous shallow shelf areas that are divided into several intrashelf basins; we assume that one of them is located north of the ramp system studied here. Based on various regional biostratigraphic zonations (e.g. Kuss & Malchus 1989; Kuss & Schlagintweit 1989; supplemented by new data from M. A. A. Bassiouni & M. Simmons, pers. comm.) the lithofacies-microfacies investigations enabled us to reconstruct the controlling mechanisms of the interdigitating continental-marine units within the Cretaceous successions. Furthermore, the shifts of the shoreline were interpreted as reflecting major onlap-offlap patterns within the stratigraphic successions of the different areas (Kuss 1992; B a c h m a n n et al. 1996). They allow the

Fig. 2. Palaeogeographical maps to illustrate the Late Aptian and Late Cenomanian environments of deposition (after Kuss & Bachmann 1996). The carbonate ramp of the northern Sinai passed laterally into a carbonate platform with a steeply inclined slope further northeast (Sass & Bein 1982). Explanation of ornaments: 1, exposed swell; 2, terrestrial realms marked by fluvial sandstones or nondeposition; 3, deltaic sediments; 4, inner platform-ramp with siliciclastic deposits; 5, inner platform-ramp with mixed carbonate-siliciclastic deposits; 6, inner platform-ramp with carbonate deposits; 7, inner ramp with carbonate shoals (northern Sinai); 8, platform margin with reefs (northern Israel); 9, platform slope-mid- and outer ramp; 10, basin; 11, position of the sections.


256


estimation of the second-order relative sea-level fluctuations (Fig. 2) that form the basis of the detailed interpretations of third-order sea-level changes described here.

Methods and basic assumptions We combined detailed field observations of different macroscopic facies types with petrographic studies, mainly of the carbonates, to interpret them in terms of depositional environments. Sedimentary structures, depositional geometries, and fossil contents are the most important macroscopic proxies. The microscopic data are based on 500 thin sections used for biostratigraphy, microfacies subdivisions, and a semi-quantitative component analysis. Considering the distribution of 24 microfacies types (MF types) within the sections, it is possible to define facies zones and thus to interpret the area of deposition for each MF type. The lateral and vertical reconstruction variability of facies allowed estimates of environmental changes during deposition and the reconstruction of a curve of relative sea-level changes for the central region (Figs 3 and 4). Although most are connected to changes of relative water depths and reflect the increasing or decreasing extents of accommodation space across the area, palaeobathymetric changes are less well evidenced, as would be expected by the reconstructed sea-level oscillations. During transgressive intervals, carbonate production and accommodation space increased and are visible in frequent shoals. Regressions can be seen in progradations of peritidal evironments. It is only during emergence of the ramp that palaeobathymetric changes caused by oscillations of the relative sea level are evidenced by karstification, rhizoliths and other features. Among the various surface types observed, some may indicate shallowing or flooding surfaces, whereas others show firmgrounds, hardgrounds or subaerial exposures. Their lateral continuation, the observed facies variations

above and below these boundaries, and the reconstructed depositional geometries allowed the interpretation in terms of sequence boundaries (SB), transgressive surfaces (ts) or maximum flooding surfaces (mfs). After defining depositional sequences (sequence numbers are related to the underlying sequence boundaries), they have been correlated first within the central area (sections Gebel Mansoura (M, ME), G. Raghawi (R) and G. Mistan (Mi), Figs 1, 3 and 4, including a few other sections not figured here). Here, we observed an excellent lateral control of sequence boundaries, systems tracts, and stacking of macrofacies patterns. The sequence stratigraphic interpretation defined here has then been extended further north (section Rizan Aneiza (RN), Figs 3 and 4) and to the south and southeast (sections Gebel Halal (H), G. Minshera (GM) and Areif el Naqa (AN 1, 2), Figs 1 and 5) within a multi-biostratigraphic frame and the progradational, aggradational or retrogradational characters of lateral facies changes. Subsequently, we studied the internal factors controlling sedimentation processes within different environments and within the different systems tracts by means of high-resolution microfacies analysis based on a semiquantitative database. As a result, we have to conclude that different interacting internal factors control sedimentation processes and their distribution within the systems tracts, among which, however, relative oscillations of sea level have been regarded as of major importance.

Stratigraphic framework Pre-Aptian deposits are represented by continental sandstones. Marine Upper Aptian sediments occur only in the northernmost Sinai (sections RN and R, Figs 2 and 3), whereas marine Albian to Cenomanian strata cover the northern and the southern regions. Lower to Upper Albian successions can be found in the central area (Fig 3 and 4), and contrast with substantially thinner Upper Albian units further

Fig. 3. Stratigraphic subdivision and correlation of the Upper Aptian to Lower Albian succession of the northern sections (for locations, see Fig. 1). The sequence stratigraphic interpretation of the northern Sinai ramp deposits is based on multibiostratigraphic data (numbers right of the sections) and the determination of the different environments of deposition, indicated in the columns left of the sections. Late Aptian ostracods (marked by rhombs) are Rehacythereis btaterensis btaterensis (1), Hechticythere croutesensis (2), Physocythere nobilis (3), Centrocythere sanninensis (morphotype B) (4), Rehacythereis btaterensis interstincta (5) and Rehacythereis phoenissa praeserotina (6). Late Aptian orbitolines (marked by squares) are Mesorbitolina (M.) subconcava (1) and Mesorbitolina (M.) texana (2). The Early Albian ostracods (marked by rhombs) are Rehacythereis phoenissa phoenissa (7), R. zuomoffeni diuturma (8) and Centrocythere sanninensis (morphotype A) (9). Early Albian ammonites (marked by circled (1)) are Eogaudryceras (E.) vocontianum, Tetragonites timotheanus rectangularis, Beudanticeras cf. beudanti, Knemiceras cf. uhligi subcompressum, Knemiceras sp. ex gr. uhligi and Knemiceras sp. ex gr. syriacum.

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M. B A C H M A N N & J. KUSS


MIDDLE CRETACEOUS SINAI RAMP south (Fig. 5). A similar decrease in thickness and stratigraphic range is true for the Cenomanian strata in the south (Fig. 5). This general distribution pattern is due to a second-order southward transgression during Mid-Albian to Cenomanian times, flooding the former continental regions (Fig. 2; Kuss & Bachmann 1996). Because of the prevailing shallow marine Aptian-Cenomanian deposits in the northern Sinai, the biostratigraphic interpretation results mainly from benthic organisms supplemented by biostratigraphic data for a few planktonic foraminifera in the upper part of the eastern sections, and ammonites from the central and eastern regions (Geyer et al. 1998). The analysis of ostracod assemblages in comparison with range charts described from age-equivalent strata in Israel (Rosenfeld & Raab 1974, 1984) allows subdivision of the Upper Aptian to Cenomanian succession (M. A. A. Bassiouni, pers. comm.). In the central and northern parts, orbitolines and dasycladacean algae were used for the biostratigraphic subdivision of the Upper Aptian to Cenomanian sediments (Kuss & Schlagintweit 1989; Kuss 1994; M. Simmons, pets. comm.). The generally accepted stratigraphic concept of the region (compare Jenkins 1990; Lewy 1990) was recently complicated by a new interpretation given by Ayyad et al. (1996) for the Cretaceous succession of Areif el Naqa (AN 1 and 2, Figs 1 and 5). Those workers described 'Cenomanian' carbonates that are, however, of Late Albian age based on an earlier biostratigraphic subdivision (Bartov et al. 1980) and new ammonite findings (Geyer et al. 1998). Moreover, the abundance of planktonic Foraminifera from the 'Albian' and 'Cenomanian' parts of the section, as described by Ayyad et al. (1996), cannot be confirmed with our data based on detailed sampling. A more detailed discussion of these stratigraphic discrepancies was given by Ltining et al. (1998). According to the generally accepted regional stratigraphic framework and the lithology expressed within all sections, two depositional units are distinguished: a siliciclastically influenced lower unit A of Late Aptian to Early

259

Albian age and an upper carbonate unit B of Mid-Albian to Cenomanian age (Figs 3-5). The Upper Aptian-Albian units correspond to the Rizan Aneiza Formation defined for the Barremian to Albian strata of Sinai (Said 1971), or the Hatira Formation as defined by Bartov et al. (1980). The Cenomanian strata correspond to the Halal Formation (Said 1971) or the Hazera Formation (Upper Albian-Cenomanian units) of Bartov et al. (1980). Unit A

The Lower Cretaceous marine succession starts with alternations of siliciclastic rocks and carbonates deposited in the northern region: sections RN, M, MiS, R, M, and ME (Figs 1, 3 and 6a). The maximum thickness of 260 m is reached in section R, which is characterized by high amounts of siliciclastic material. In general, unit A exhibits limestone-marl-clay alternations with common siltstone and sandstone intercalations, because of the development of a delta system in the central region (Fig. 6a; section 4.1). The limestones range from pure mudstones to oolithic grainstones with high contents of siliciclastic material (mainly quartz, subordinate glauconite and limonite). Ferruginous ooids are common in all sections and are well dispersed in the delta-influenced clay and marl beds or enriched in sandstone layers. Ferruginous crusts are locally common, often in combination with sandstone intercalations. The lower part of unit A is of Late Aptian age (Arkin et al. 1975), as is confirmed by ostracod assemblages that allow subdivision of the Upper Aptian succession into two intervals (M. A. A. Bassiouni, pers. comm.): the lower part has Rehacythereis btaterens& btaterens& (rhomb 1), Hechticythere croutesensis (rhomb 2), Physocythere nobilis (rhomb 3), and Centrocythere sanninensis (morphotype B) (rhomb 4), which was found in section R only (Fig. 3), whereas Rehacythereis btaterensis interstincta (rhomb 5) and Rehacythereis phoenissa praeserotina (rhomb 6) occur in the upper Upper Aptian strata of sections R and RN (Fig. 3). The

Fig. 4. Stratigraphic subdivision and correlation of the Middle Albian to Cenomanian succession of the northern sections (for locations, see Fig. 1). The sequence stratigraphic interpretation of the northern Sinai ramp deposits is based on multibiostratigraphic data (numbers right of the sections) and the determination of the different environments of deposition, indicated in the columns left of the sections. The southward adjoining sedimentary succession is shown in Fig. 4 (for legend, see Fig. 3). Besides the Albian ostracod Centrocythere sanninensis (morphotype A) (rhomb 9), the ammonite Knemiceras sp. ex gr. uhligi (circle 2), several orbitolines (marked by squares) are of stratigraphic importance: different evolutionary stages of Mesorbitolina (M.) texana (1) and Mesorbitolina (M.) subconcava (2) occur in several horizons with Palorbitolina heddini (3). Conicorbitolina conica (4) indicates an Early Cenomanian age.




MIDDLE CRETACEOUS SINAI RAMP A p t i a n - A l b i a n boundary cannot be clearly defined; however, the lower part of the Lower Albian sediments contains Mesorbitolina (M.) subconcava (square 1) and early evolutionary stages of Mesorbitolina (M.) texana (square 2; M. Simmons, pers. comm., Fig. 3), gradually evolving in the upper parts, where they occur together with Palorbitolina heddini (square 3; section RN, Fig. 4). Ostracods of Rehacythereis phoenissa phoenissa (rhomb 7) and R. zuomoffeni diuturma (rhomb 8; M. A. A. Bassiouni, pers. comm.) were found in the Lower Albian part of section R (Fig. 3) together with Centrocythere sanninensis (morphotype A) (rhomb 9); the latter, however, occurs up to the Upper Albian strata (section GM, Fig. 5). Further stratigraphic data come from two ammonitebearing horizons of sections M and ME (Figs 3 and 4), indicating an Early to Mid-Albian age (Geyer et al. 1998). The first lies in sequence 2 and holds a rich fauna of Eogaudryceras (E.) vocontianum, Tetragonites timotheanus rectangularis, Beudanticeras cf. beudanti, Knemiceras cf. uhligi subcompressum, Knemiceras sp. ex gr. uhligi, and Knemiceras sp. ex gr. syriacum (circle 1; Fig. 3), whereas the second within sequence 4 holds only Knemiceras sp. ex. gr. uhligi (circle 2; Fig. 4).

Unit B Unit B can be subdivided into a Middle-Upper Albian part and a Cenomanian part. Sediments of the former were identified from seven sections in the studied region: sections Mi, MiS, R, M, GM, H, and AN 2 (Figs 4 and 5). The maximum thickness of c. 160 m is reached within the northern section Mi. In the central to northern region, the Middle-Upper Albian succession is dominated by limestone-marl alternations. The input of coarser siliciclastic material and ferruginous ooids ended, owing to a southward retrogradation of the shoreline (sections Mi, MiS, R and M, Figs 2 and 4). Oolithic and bioclastic sediments alternating with rudist biostromes, as well as mudstones, characterize the carbonates. Within the southern area (sections GM, H and AN, Fig.

261

5), the siliciclastic input is still represented by some sandstone and siltstone layers, owing to a nearshore position. The abundance of highenergy grainstones within the limestones evidently decreases in the south coevally with an increase in lower-energy wackestones. Orbitolina assemblages of stratigraphic value occur in several layers of unit B in the northern sections (Fig. 4), comprising different evolutionary stages of Mid-Late Albian Mesorbitolina (M.) subconcava (square 1) together with Mesorbitolina (M.) texana (square 2; M. Simmons, pers. comm.). Late Albian ostracods are 'Ahmadiura' bisulcata (rhomb 10) and Curfsina nuda azraqaensis (rhomb 11; M. A. A. Bassiouni, pers. comm.) from the southern section GM (Fig. 5) co-occurring with Centrocythere sanninensis (morphotype A) (rhomb 9). Ammonites of Knemiceras uhligi subcompressum found in marls of section AN 1 (circle 3; Fig. 5) confirm the Late Albian age of these units (Geyer et aL 1998). Cenomanian sediments are present within the entire region. Because of later erosion, the northern sections (Mi and M) include only the lowest Cenomanian, whereas two sections in the south (GM and AN) document complete Cenomanian successions, reaching a maximum thickness of 140 m (section GM, Fig. 5). The Lower Cenomanian succession is represented by limestones and dolomites with some marlstone intercalations in the north, the latter becoming more frequent in the Cenomanian strata of the southern areas. Within the Upper Cenomanian succession, the siliciclastic component decreases in the southern sections as well, reflecting the southward shifting coastline to southern Sinai areas as a result of the second-order sea-level rise (Kuss & Bachmann 1996). Rudist biostromes are intercalated in the Lower Cenomanian strata of the northern sections (M and Mi, Figs 4 and 6c) and decrease in frequency towards the south. Strong dolomitization of the Lower Cenomanian carbonates complicates the exact determination of the A l b i a n - C e n o m a n i a n boundary in the north, defined by the first

Fig. 5. Stratigraphic subdivision and correlation of the Upper Albian to Cenomanian succession of the southern sections G. Halal, G. Minshera, and G. Areif el Naqa (for locations, see Fig. 1). The sequence stratigraphic interpretation of the ramp deposits and the different environments of deposition are indicated (for legend, see Fig. 3). The northward adjoining sedimentary succession is shown in Fig. 4. Index-fossils of the Upper Albian succession comprise ostracods (marked by rhombs) Centrocythere sanninensis (morphotype A) (9), 'Ahmadiura' bisulcata (10), Curfsina nuda azraqaensis (11), and ammonites Knemiceras uhligi subcompressum (marked by circle 3). The Cenomanian units include foraminifera Praealveolina cretacea (marked by *), and ostracods (marked by rhombs) Veeniacythereis maghrebensis (12), Phlyctocythere citreum (13), Procytherina cuneata (14) and Perissocytheridea sohni (15).



Fig. 6. (a) Section G. Mansoura (M) of the central area illustrates the two ramp settings: the lower foothill areas with delta-dominated Upper Aptian-Lower Albian siliciclastic deposits (unit A) and Middle Albian-Lower Cenomanian massive limestones (unit B) at the top (total thickness 340 m). (b) Section G. Minshera (GM) of the southern area represents cyclic sedimentation patterns deposited in shallow marine to tidal flat environments of unit B. (c) Small rudist biostrome (with Eoradiolites lyratus, J.-P. Masse, pets. comm.) of section M within the late HST of sequence 8. (d) Oolitic shoal with large-scale cross-bedding within the uppermost TST of sequence 9 at the G. Mansoura (M). (e) Rhizoliths within mudstones at the sequence boundary 11 at the G. Mistan (Mi). (f) The sharp erosive contact at the top of a rudist biostrome (Sellaea facies, J.-P. Masse, pers. comm.) marks a higher-frequency cycle boundary at the G. Rhagawi (R) within the third-order HST of sequence 8.


MIDDLE CRETACEOUS SINAI RAMP (M. occurrences of Conicorbitolina con 9 Simmons, pers. comm.) in section M (square 4; Fig. 4). At section AN 2, further south (Fig. 5), the marly En Yorqeam Member defines the Albian-Cenomanian boundary (Lewy 1990), confirmed by the first occurrences of Foraminifera Chrysalid 9 gradata and Pseudedomia drorimensis in thin limestones above the ts of sequence 11 (Fig. 5). Moreover, the first occurrences of Praealveolina cretacea (*) in the transgressive systems tract (TST) limestones above favour an Early Cenomanian age and can be compared with findings in section GM (Fig. 5). Ostracods of Veeniacythereis maghrebensis (rhomb 12) and Perissocytheridea sohni (rhomb 15; section AN 2), and Phlyctocythere citreum (rhomb 13) and Procytherina cuneata (rhomb

263

14) from section H (Fig. 5) confirm the Cenoman 9 age (M. A. A. Bassiouni, pers. comm.). The Cenomanian-Turonian boundary is clearly defined by deep-water limestones with planktonic Foraminifera and a characteristic association of earliest Turonian ammonites, described by Bartov et al. (1980). Facies patterns

of the

ramp

Within the limestones and marls, 24 microfacies types (MF type) were distinguished according to their texture, g r o u n d 9 and components (Table 1). For each MF type, the environment of deposition has been interpreted on the basis of comparisons with descriptions from corresponding settings of Recent (Purser & Evans 1973)

Table 1. Classification of the Late Apt9

to Cenomanian MF types of the northern Sinai carbonate ramp defined for the carbonates of units A and B

texture onents

MF-type

1 2 3 4 5 6 7 8 9 10 ii i2 13 14 15 16 17 18 19 20

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9 9 [] 20-30 . 300 m thick shallow marine succession of Albian-Lower Cenomanian carbonates. We interpret the rather gradual lateral transitions of thinning and thickening rudist biostromes, indicated by different magnitudes from several localities in northern Sinai, as related to the ramp geometry. No shelf margin break can be found in these units, which formed along gently north-dipping slopes. Abundant high-energy shoals with common channel structures filled with grainstones, and large extended tidal flat


MIDDLE CRETACEOUS SINAI RAMP areas (see below) contribute to the carbonate succession and lead to the classification of the northern Sinai ramp as strongly wave and tide influenced (sensu Burchette & Wright 1992). The following descriptions refer to two simplified ramp models (Fig. 7a and b); more specific and detailed characteristics of the sedimentation processes will be given below.

The Upper Aptian to Lower Albian ramp; unit A The Upper Aptian to Lower Albian deposits of the shallow ramp are restricted to the northernmost Sinai, from the Maghara area to the north, whereas south of this region, continental sandstones prevail (Figs 2 and 7a). The carbonates of unit A are characterized by the occurrence of various MF types, containing a high amount of siliciclastic material (MF types 4, 5 and 15). These carbonates are intercalated with pure claystones, siltstones and sandstones (all with changing contents of iron ooids, Table 1, Figs 3 and 4), which were interpreted as delta sediments. Although not studied in detail, field observations and thin-section data allow us to ascribe them to sedimentation processes of the delta plain, with its meandering floodplains, swamps, and beach complexes. Sediments of the steeper delta front and the broadly sloping prodelta could not be

265

recognized. Sedimentation processes on the shallow north-dipping ramp were markedly influenced by this delta regime, comparable with a tide- and wave-dominated delta system, formed along a 'shoal water' profile (Postma 1990). Considering the lateral distribution of the siliciclastic intercalations within the U p p e r Aptian-Lower Albian unit A (Figs 3 and 4), two regional trends are visible. A vertical decrease of siliciclastic intercalations is obvious and attributed to the termination of deltaic-influenced sedimentation in unit B (see below). Laterally, especially in the lower sequences of unit A, a general decrease in grain size from south to north can be observed, which is more conspicuous in sections from localities (not described) a few kilometres further south of sections M and ME (see Fig. 1). There, almost pure sandstone units with conspicuously overlapping channel casts occur and indicate floodplain deposition, with only minor imprints of marine ingressions. Most of the delta sediments further north (central area between sections M and Mi, Figs 1 and 3) are composed of sandstones, fine siltstones and claystones that interfinger laterally and vertically with marls and carbonate lithologies. Whereas the silts and clays settled out of suspension in the delta-front or lagoonal complexes further north (sections R and RN), sandy wedges and tongues are scattered through all

Fig. 7. Late Aptian to Cenomanian ramp geometries and facies distribution of the northern Sinai. The following facies zones can be distinguished within units A and B: 1, terrestrial sediments; 2, deltaic sediments: claystones and sandstones (MF types 5); 3, peritidal environments: mudstones (MF types 20, 21 and 22), locally with siliciclastic input; 4, protected inner ramp with siliciclastic input: siltstones, marlstones and wackestones to packstone (MF types 20 and 24) and protected inner ramp without siliciclastic input: marlstones, wackestones and packstones (MF types 16-20, 22); 5, open marine shallow ramp: packstones (MF types 8 and 10-14); 6, open marine shallow ramp with siliciclastic input, silty and sandy carbonates (MF type 15); 7, oolitic and bioclastic shoals: grainstones (MF types 1, 3, 6 and 7), locally with siliciclastic input (MF types 2, 4 and 5); 8, rudist biostromes, packstones, floatstones and bindstones (MF types 7, 8 and 9); 9, mid-ramp (MF type 23) (for classification of MF types, see Table 1).


266


fine-grained lithologies (including the carbonates) even in the southern sections of G. Mansoura (Figs 3 and 6a). They tend to build narrow lenses of sand that are abundantly cross-bedded, with abundant ripple cross-lamination, scourand-fill structures and discontinuous clay lenses. These sedimentary structures characterize various distributary channel systems of the river (Davis 1983). Besides microscopically identified coal particles, large pieces of driftwood occur, even in complexly cross-stratified sandstones of the northernmost section RN (Fig. 3, below SB 1), which may represent distributary mouth bar sands. Pelitic lithologies of finely laminated siltstones and claystones occur in increased numbers and thicknesses in the central and northern sections, and are interpreted as wide shallow interdistributary bays and marshes. Moreover, clayey intercalations (a few centimetres to 80 cm thick) with varying amounts of ferruginous ooids, root casts, and burrows are frequent. A marine origin of these iron ooids is indicated by several nuclei composed of shells of benthic organisms. The dominance of terrestrial or degraded marine kerogens and high kaolinite contents confirm a strong input of land-derived particles in these pelitic deposits. Several clayey beds suggest a mixing of river water and marine water, indicated by low diverse microfaunas, especially of smaller benthic Foraminifera. Most of the deltaic siliciclastic intercalations in unit A with disconformable contacts overlie carbonates of the late highstand systems tract (HST) (Fig. 3). As a consequence, we attributed periods of delta progradation in our model (Fig. 7a) to lowstand conditions. Although geometric considerations suggest that enough accommodation space can be expected mainly in the northern sections, we also identified several deltaic units in the central area, where they were deposited during late lowstand systems tract (LST), when a gradually rise in relative sea level commenced and created new accommodation space. In numerous examples, deltaic siliciclastic material (up to several metres thick) is overlain by marls or limestones of the next TST, which often holds extraclasts of reworked and redeposited unlithified siliciclastic deposits. These abrupt lithological changes indicate that the delta progradation terminated several times, especially during rises of relative sea level. Laterally, the deltaic units often interfinger towards the north with marls, partly enriched in oysters, gastropods and locally in orbitolines (MF type 24), which have been attributed to a restricted inner ramp facies belt. Wackestones and packstones, often with abundant pseudopeloids, benthic foraminifera, oncoids,

bioclasts and intraclasts (MF types 11-14) indicate a low-energy, shallow ramp facies belt with sporadic and subordinate siliciclastic input (MF type 15, Fig. 7a). Rare and small rudist biostromes developed in these backshoal areas with lateral extensions of a few square metres (sections RN, M and Mi, Figs 3 and 6c) and are described below. High-energy conditions in the distal inner ramp facies belt are indicated by the deposition of carbonate shoals composed of large-scale cross-bedded bioclastic and oolithic grainstones (MF types 1-3), which are often enriched in quartz sands, iron ooids and ironimpregnated lithoclasts (MF type 4). The strong siliciclastic input of the delta system still influenced all shallow ramp environments, with the increasing and decreasing rates of siliciclastic material being controlled by progradation and retrogradation of the Late Aptian-Early Albian shoreline. Mid-ramp sediments were observed only in the northernmost section RN (Fig. 4); they are similar to those of unit B and are therefore described in the next section.

The Middle Albian to Cenomanian ramp; unit B During Mid-Albian to Cenomanian times the southward extension of the ramp increased, owing to a second-order rise of sea level, which resulted in a contemporaneous flooding of the former delta system and retrogradation of the shoreline (Fig. 2). As a consequence, a major decrease in siliciclastic input took place, and marine sediments also occur within the southern exposed sections GM, H, and AN 1, 2 (Figs 4, 5 and 7b). Thus, the entire ramp was characterized by predominance of carbonates. The Middle Albian to Cenomanian inner ramp has been subdivided into five main facies belts, which in general strike east-west. They are described here from south to north, i.e. from proximal to distal ramp settings comprising peritidal environments, restricted inner ramp, shallow ramp with rudist biostromes, shoals, and midramp environment (Fig. 7b). (1) Extensive peritidal environments are developed several times within unit B. In the northern region (sections Mi, MiS, R and M, Fig. 4) they are commonly formed by alternations of dolomites (MF type 22) with mudstones containing bird's-eyes (fenestrae) and microbial laminites (MF type 21). Moreover, pure dolomitic marls and mudstones to wackestones with few bioclasts (ostracods, gastropods, benthic Foraminifera), and few ooids or intraclasts (mudclasts or reworked microbial mats)


MIDDLE CRETACEOUS SINAI RAMP are frequent. These successions, of a few metres to 30 m thickness, show lateral extensions of hundreds of metres (Fig. 6a), and their reconstructed extensions cover several kilometres (e.g. late HST of sequence 7, Figs 4 and 8). Further to the south, a peritidal subfacies characterized by siliciclastic admixtures, alternations of pure dolomites (MF type 22) and dolomitic marls occurs. Intercalations of wackestones (MF type 15) and marls enriched in quartz silt are typical (sections GM, AN 1 and AN 2, Fig. 5). Rhizolithic layers (Fig. 6e), tepee structures or mud cracks may interrupt the deposition and reflect repeated episodes of emergence. (2) Sediments of protected inner ramp environments are common within the entire region. Miliolid wackestones (MF type 17) with varying amounts of small benthic Foraminifera (e.g. textulariids, praealveolinids and cisalveolinids in the southern sections), gastropods and ostracods, as well as calcareous debris, were observed within most sections, but more frequently in the south. The same is true for bioclastic wackestones and packstones with diverse dasycladaceans and udoteaceans (MF type 16), benthic Foraminifera (e.g. textulariids, miliolids, locally praealveolinids and cisalveolinids), bioclasts and locally pseudopeloids or lithoclasts. Within all sections these two MF types form alternations, up to a few metres thick, with marls, often containing oysters and gastropods (MF type 24), or may overlie rudist floatstones. Restricted environments are reflected by mudstones comprising only minor shell debris and lithoclasts (MF type 20), and wackestones to packstones with abundant ostracods and only subordinate bioclasts (MF type 18). The latter are limited to the southeastern part of the region (sections AN 1, AN 2, GM and H, Fig. 5). A few Cenomanian ostracod assemblages in section AN 2 clearly indicate brackish conditions (M. A. A. Bassiouni, pers. comm.). Because of the brackish influence and the minor siliciclastic deposition here as against sections H and GM (Fig. 5), a narrow bay is assumed to have developed in this area, resulting in extended lagoonal areas during times with decreasing fresh-water influence. As extensive barriers were not observed within the southern outcrops, protected to restricted inner ramp environments in most sections may have resulted from wide extended shallow-water areas. (3) Open marine shallow water faunal associations imply a shallow ramp facies belt without major influences of wave activity, characterized by alternations of wackestones and packstones that commonly comprise fine-grained bioclastic varieties with benthic foraminifera (such as

267

textulariids and lituolids, Cuneolina sp., Chrysalidina sp. and miliolids), few bivalves and dasyc|adaceans, and varying amounts of pseudopeloids, calcareous debris and lithoclasts (MF type 14). This MF type indicates the transition to the protected areas of the inner ramp in the southern sections (AN 1, AN 2, H and GM, Fig. 5). Bioclastic wackestones and packstones with quartz (MF type 15) occur, however, with less frequency. Fine-grained to coarser packstones with intraclasts (MF type 13) or pseudopeloids (MF type 12), often enriched in echinoderms, indicate depositional areas with moderate to low energy within the entire region. Coarse-grained packstones with varying amounts of diverse shallow marine biota (such as small and coarse agglutinated foraminifera, orbitolines, cyanophyceans, dasycladaceans, udoteaceans, bivalves, bryozoans, gastropods, and fragments of rudists with encrusting calcareous sponges), oncoids as well as lithoclasts (MF type 11), characterize the shallow ramp facies belt of the northern sections (RN, Mi, MiS, R and M, Fig. 1). Orbitolines are common within nearly all shallow ramp MF types of the northern area, with highest frequencies in marls and bioclastic packstones. Nodular and bioturbated limestones of MF type 10 commonly contain ooids, derived from nearby shoals. These units may reach 20 m thickness, including some marly beds that contain similar components. In contrast to the older ramp system of unit A, rudist biostromes are common and alternate with protected and shallow ramp deposits, and in a few cases also with oolithic grainstones (Figs 6f and 8). Two major groups of rudists can be differentiated within the Upper Aptian-Lower Cenomanian carbonates of north Sinai (J.-P. Masse, pers. comm.; Bartov et al. 1980). Caprinids of the 'Sellaea Facies' occur in large biostromes of Late Albian-Early Cenomanian age (e.g.G. Raghawi, early HST of sequence 9, Figs 4 and 6f). These stacked recumbent caprinids form massive (2.5-9 m thick) frameworks of monotypic associations that may exhibit nearly constant thicknesses and faunal compositions over >1 km laterally. Radiolitids occur in several smaller Upper Aptian to Albian biostromes (e.g. the Late Albian Eoradiolites lyratus of section M, within late HST limestones of sequence 8, and of section AN 1, within late HST limestones of the same sequence). These 15-40 cm high congregations of clustered elevators (Fig. 6c) grew constratally without any supporting biogenic frameworks (Gili et al. 1995) and can be followed laterally over several tens of metres.


268


Both groups of rudists were found in units that display rather gradual lateral thinning and thickening, and that exhibit no relief (with no forereef facies), even in large-scale correlations, where profiles run normal to the palaeo-coastline along the gentle slopes of the ramp system (Fig. 4). A medium-scale rudist lithosome (sensu Skelton et al. 1995) is exposed at section RN (directly underlying SB 1, Fig. 3), and represents a more general example of the rudist deposition in Sinai. Here, the radiolitid lithosome reaches a maximum thickness of 1.60 m thinning to 0 m over a 65 m lateral distance. Whereas the thicker parts are composed of >85 % rudists, their total content decreases towards the flanks with a coeval increase in oncoid-bearing lithologies. We conclude that similar lateral transitions have to be assumed for many of the larger and smaller rudist biostromes. They may have formed slight elevations of great lateral extent on the sea floor, which resulted in shallow undulating bottom topographies. As a consequence, these rudist biostromes might have influenced the bottom water currents and thus the establishment of local protected areas. (4) Carbonates of the shoal facies are widespread in the central areas, less frequent in the northern areas, and absent in the southern sections. They form massive beds with thicknesses of several decimetres, often up to several metres thick, and always characterized by sinuously undulating bed surfaces that exhibit lateral thickening and thinning. Very conspicuous lowangle cross-bedding is visible macroscopically (with decimetre-scale foresets), as well as in smaller scales within single foresets (Fig. 6d). Three MF types are most frequent oolithic grainstones (MF type 1) with up to 90% of ooids, bioclastic grainstones with ooids (MF type 3) and bioclastic grainstones (MF type 6). The last two are composed of well-rounded fragments of echinoderms, bivalves, orbitolines, dasycladaceans, and udoteaceans. Further constituents of minor importance are textulariids, lituolitids, gastropods and bryozoans. Plastically deformed lithoclasts of unlithified and lithified mudstones and ooid packstones may be locally common. These MF types can be compared with grainstones of similar textures and spatial distributions, described from several modern environments such as the Arabian Gulf (Loreau & Purser 1973). Oolitic and bioclastic grainstones were deposited in high-energy barrier shoal complexes separating outer ramp deposits from protected inner ramp areas and peritidal environments. The predominance of ooids or bioclasts may have resulted from oscillations of

the relative sea level and thus changes of the accommodation space (see discussion). (5) Mid-ramp sediments are represented by few examples of marls alternating with thin limestones of the northernmost sections RN and M (Fig. 4), and by Upper Cenomanian-Lower Turonian marls in the southern section AN 2 (MF type 23). Mid-ramp limestones occur in sequence 6 of section RN and change laterally into shallow marine carbonate units of the inner ramp. They are composed of wackestones and mudstones with planktonic foraminifera, together with changing amounts of smaller benthic foraminifera, spiculae, echinoderms, pseudopeloids, and mudclasts. They cannot be followed laterally over longer distances, and seem to increase in frequency and thickness further north, as described from subsurface sections (Ayyad & Darwish 1996).

Sequence stratigraphic interpretation The following sequence stratigraphic interpretation of the Upper Aptian to Cenomanian succession, and the reconstruction of the local relative sea-level changes, follows the nomenclature of Van Wagoner et al. (1988), Vail et al. (1991), and Sarg (1988), supplemented by modern concepts especially described for sequence stratigraphic interpretations of carbonates (e.g. Handford & Loucks 1993; Tucker et al. 1993; Read 1995; Wright & Burchette 1996). Although parasequences are often observed, this paper is focused on the third-order sequences only. Owing to the different major depositional environments, sequence stratigraphic patterns vary in these Upper Aptian-Lower Albian (unit A) and Middle Albian-Cenomanian (unit B) successions (Figs 3-5). The regional distribution and shifting of facies belts is illustrated within three facies maps (Fig. 8). They reflect three different palaeogeographical situations during sequence 7, which are due to different relative sea levels. As sequence 7 crosses the boundary between unit A and B (Fig. 4), sedimentological characteristics of both units are visible in the respective LST, TST and HST. U p p e r A p t i a n to L o w e r A l b i a n ; unit A In the central area, the major depositional characteristics of unit A are summarized on the basis of five sections (Fig. 4); another two sections in the southern area (Fig. 5) comprise the uppermost part of unit A only. Therefore, only the northern sections were taken into consideration in the following; they reflect a 30-40 km wide segment of the shallow, delta-influenced


MIDDLE CRETACEOUS SINAI RAMP

269

Fig. 8. Palaeogeographical maps representing three different situations during sequence 7 in the central region (Fig. 3) to illustrate the main differences in the distribution and the occurrence of the facies belts between the individual systems tracts. These maps include data of sections not discussed here. LST: Broad areas of emergence and terrestrial sediments occur in the south. Deltaic sediments covered the northern part of the area, interfingering with restricted ramp environments further north. TST: A strong retrogradation took place. The northern area was covered by a broad belt of bioclastic shoals; here protected inner ramp environments developed in protected areas behind the shoals and interfinger with siliciclastically influenced restricted inner ramp environments further south. HST: The southernmost area was covered by tidal flats. Northward protected inner ramp environments and rudist biostromes passed into shallow ramp areas with large and extended rudist biostromes. inner ramp, comprising seven sequences (bounded by SB1-SB8). The boundary to unit B lies within the seventh sequence (between SB7 and SB8; see Fig. 4). Sequence boundaries (SB) are m a r k e d by h a r d g r o u n d s and ferruginous crusts, which reflect sedimentary discontinuities in the lower part of the unit (Fig. 4). Within the upper part, sequence boundaries are recorded by rhizolithic layers in the northern distal sections (RN, Mi, MiS-SB6, SB7) and by the deposition of continental sandstones (sections M, SB6) or strong karstification (sections M and SB 7; Fig. 4) in the south, which together document the emergence of large areas of the shallow ramp and the prog r a d a t i o n of the facies belts. Because of a higher-frequency cyclicity it is sometimes not possible to indicate the exact position of the

sequence boundary: more than one emergence horizon within a short vertical distance (e.g. SB3 in section M E ) or high-frequency changes in facies may be visible. In such cases, the strongest emergence or the most significant facies change, respectively, is defined as the sequence boundary. The increase in siliciclastic material resulting from the contemporaneous northward progradation of the delta system (Figs 3 and 8) determines the LSTs within the entire unit A. As a rule, the deposition of deltaic sediments overlying pure marine strata indicates a northward progradation of the facies belts (compare section above). Furthermore, the emergence at the base of a deltaic unit (clearly visible in sequence 6 of section R) confirms the interpretation as an LST deposit. We explain the local variations of the


270


delta facies within short distances as bein~ due to prograding delta-lobes, characterized by increasing numbers of sandstones (often with channel casts) with many claystone intercalations. They also exhibit increasing contents of iron ooids, ferruginous lithoclasts and drift wood. The transgressive surface (ts) is the first significant marine flooding surface across the shallow ramp (Van Wagoner et al. 1988). On the northern Sinai ramp it is recorded by a transition from deltaic sediments to restricted inner ramp sediments in nearly all sections, associated with a sudden cease in the deposition of coarser siliciclastic material (Fig. 4). Within the lower part of unit A (Fig. 3), a gradual rise of the relative sea level is generally evident from the TSTs. Thus, a high-energy inner ramp facies is not established before the late TSTs and is more common in the northern sections (R and RN); it is well documented in sequences 1 and 2 (Fig. 8). The increasing retrogradation of the facies belts is the most important feature of the TSTs. Slightly different patterns, however, are visible in the upper parts of unit A: a stronger transgression is reflected by the establishment of open marine, high-energy conditions (characterized by cross-bedded grainstones) directly overlying the transgressive surfaces in nearly all sections (e.g. sequences 5 and 6, Fig. 4). As a consequence, the entire shallow ramp was flooded within short time intervals during TSTs. The position of the maximum flooding surface (mfs) cannot be located exactly within all sections. A change in facies mostly indicates the beginning of progradation, sometimes combined with renewed siliciclastic input (e.g. sequence 2, Fig. 3), and thus defines the base of the HSTs. The HSTs are composed of deposits originating from different facies belts: sediments of restricted inner ramp environments mostly occur in the proximal sections (ME, M or R) and carbonate shoals are frequently present within the distal sections (Mi, MiS and RN). Many small-scale facies changes are due to higher-frequency sea-level fluctuations. M i d d l e A l b i a n to Cenomanian; unit B Within unit B, the sequence stratigraphic interpretation is demonstrated in detail for the northern to central area along a N-S striking ramp transect of 40 km extension comprising five sections (Fig. 4). Here, the Middle Albian-Lower Cenomanian succession can be subdivided into six sequences (overlying SB8-SB13). Additionally, four sections further in the south were analysed and interpreted with respect to sequence stratigraphy (Fig. 5). These

sections comprise the upper part of the Albian and the Cenomanian succession, and indicate the sedimentation patterns of the sequences in the areas further south. Within the following description, the sequence stratigraphic interpretation is given first for the central area, followed by the extended interpretation including the southern area and the complete Cenomanian succession. All SBs of unit B are characterized by emergence horizons. These are indicated by rhizolithic layers (Fig. 6e) and supratidal sediments in the northern sections (Mi and MiS), and by tepee structures and palustrine limestones in the central region (R and M); strong karstification occurs only at SB 8 (Fig. 4). In intervals with more than one emergence horizon, the most distinct was used to define the SB (e.g. SB9, section R). In the southern sections, the Middle Albian to Cenomanian SBs are characterized by continental sandstone intercalations (section H, Fig. 5), by brackish or restricted sediments overlying marine strata, and by few emergence horizons (sections A N 1, AN 2 and GM). The LST sediments of unit B generally reflect deposition within the proximal inner ramp facies belts. Tidal flats and restricted inner ramp sediments are more frequent in the central region (sections M and R, Fig. 4) and especially frequent in sequence 9, where they occur in all sections. Here, the emergence horizons within the LST reflect relative sea-level fluctuations of higher frequencies within the late LST (section R). Furthermore, there are shallow open marine sediments including subordinate small biostromes (sequence 8) that are well developed in the northern sections (MiS and Mi). Rhizolithic horizons are frequent within the LSTs. Missing LST deposits within sequences 10-12 (Fig. 4) are interpreted as reflecting the emergence of the whole area during that LST. In Early Cenomanian times, the third-order cycles (sequences) were superimposed by a secondorder sea-level rise, and additional accommodation space was created. As a consequence, the LSTs are now often characterized by thick peritidal deposits (e.g. sequence 13 of section M). In the southern area, sandstones (section H), brackish sediments (section AN 2), peritidal sediments often with siliciclastic influence (section GM), and deposits of restricted lagoons (sections AN 1, GM and H, Fig. 5) commonly characterize the LSTs. Often, the transgressive surfaces cannot be exactly defined in unit B. However, they can be recognized in the transition from aggradational to progradational sedimentation patterns in some sections. In tidal-flat areas, increases in bed thickness reflect an increase in accommodation space (sequence 9, section M, Fig. 4),


MIDDLE CRETACEOUS SINAI RAMP here interpreted as a transgressive surface. In some cases an abrupt change from low-energy shallow marine environments to massive carbonate shoals shows the increasing accommodation space and thus marks the transgressive surface (sequence 8, all sections, Figs 4 and 5). Because of the absence of LST sediments within sequences 10-12, their transgressive surfaces coincide with the sequence boundaries and are marked by coastal onlap. The TSTs are nearly always characterized by the occurrence of open marine high-energy carbonates, documenting a retrogradation of facies belts, expressed by oolitic and bioclastic shoals frequently covering wide areas of the ramp. Thus, different amounts of transgression are expressed by different extensions of the areas covered by carbonate shoals (Fig. 6d). They may be restricted to the northern sections (e.g. sequence 9) or may cover the entire northern and central region (e.g. sequence 8). In all cases, these shoals interfinger with restricted inner ramp and open marine shallow ramp deposits to the south. The TSTs in the southern region are mostly reflected by lagoonal sediments, as against the more restricted sediments of the LSTs. The maximum flooding surface coincides with the transition from open marine high-energy to low-energy (lagoonal) environments in the northern and central region. The retrogradational facies patterns ended in the late TSTs and gradual progradation characterized the following HSTs, which resulted in the filling of the former proximal inner ramp areas during times of stable sea level. Moreover, extended rudist biostromes were established during the early HST within the distal inner ramp (sections H, M, R, Mi and MiS, Figs 4, 5 and 8). All biostromes alternate with extended areas of restricted inner ramp environments and shallow marine sediments, and interfinger with them towards the south, where biostromes are less common. Tidal flat progradation often started during the late HSTs; subsequently, they directly overlie the rudist biostromes (sequence 7, sections M, R and Mi; sequence 9, sequence 10, sections R and Mi). Furthermore, peritidal sediments are very common within the entire Cenomanian HSTs in the south.

Microfacies and sequence stratigraphy; a detailed depositionai model for the Middle Albian and Cenomanian units The distribution of MF types within different systems tracts was analysed for the sections RN, Mi, R, M, GM, H, A N 1 and A N 2; this means

271

that the abundance of the individual MF types within LST, TST and HST has been quantified for both units A and B (Table 2). A total of 280 thin sections (211 within unit B) out of about 500 have been considered. The others were excluded to avoid a wrong weighting of frequently sampled MF types. Thus each MF type was only counted once per systems tract per sequence within each section. Because an equidistant sampling was not possible, the resulting distribution gives only an impression of the quantitative MF type distribution. It has to be mentioned that, within strongly dolomitized parts of the sections, only a few samples were taken, and thus the often dolomitized rudist biostromes and the tidal flats, for example, are not well represented. MF types 9,19, 23 and 24 (Tables 1 and 2) are not considered here because of their rareness within the parts of the succession that were interpreted with respect to sequence stratigraphy. For the same reason, the distribution of MF types 2, 5, 7,10 and 15 is not very significant. However, the results show clearly that the occurrence of most MF types is correlated with the development of the individual systems tracts and thus related to relative sea-level changes. Additionally, the abundance of the MF types within the individual sections is considered semi-quantitatively for the different systems tracts (Fig. 9), to obtain an impression of the lateral distribution of the MF types. A detailed model for the different systems tracts can be developed on the basis of the above data as well as the macrofacies and sedimentary structures (Fig. 10). As mentioned above, the microfacies database is more extensive for unit B; thus the detailed model is given for the Middle Albian to Cenomanian succession.

Lowstand systems tract Because of the frequent emergence of the inner ramp during most LSTs, MF types representing LST deposits are not very common. However, some MF types are characteristic of LST sediments (Table 2), as follows. Within LSTs of unit A, grainstones with terrigenous components (MF type 4) and sandstones (MF type 5) are the most common MF types. Additionally, marlstones, claystones and sandstones frequently occur in LSTs of the northern sections (Figs 3 and 4) reflecting an increasing deltaic influence during the Late Aptian-Early Albian lowstands of sea level. Within unit B three groups of MF types are common in the LSTs (Table 2, Fig. 9). (1) Mudstones and dolomites (MF types 21 and 22) associated with peritidal carbonates were found in nearly all sections along the transect. Within section GM, these dolomites are


272


Table 2. Distribution of M F tvpes within the different systems tracts; results from sections RN. R. M. MI, H, GM

and A N (for location, see Fig. 1).

Percentage of thin sections

Number of thin sections

Microfacies type i 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24

Grainstone with ooids Grainstone with intraclasts and radial ooids Bioclastic grainstone with ooids Grainstone with ooids and terrigenous components Sandstone with sparitic matrix Bioclastic grainstone with intraclasts Bioclastic packstone to grainstone with rudist debris Floatstone with rudists and crusts Floatstone to packstone with encrusting organism Bioclastic packstone with ooids Bioclastic packstone (with intraclasts) Peloidal wackestone to packstone with bioclasts Packstone with intraclasts Peloidal wackestone to packstone rich in Foraminifera Bioclastic wackestone to packstone with quartz Bioclastic wackestone to packstone with dasycladaceans Mudstone to packstone with abundant miliolids Wackestone to packstone with ostracods Wackestone with few ooids Mudstone Mudstone with bird's-eyes Dolomite Mudstone with planktonic Foraminifera Marls with bioclasts Total

~

~

~

~

~

~

~

~

10 5 10

2 4 6

8 1 4

0 1 1

7

3 5

3

0

1 4

20 10

70 60 50

30 20 40

10 5 13

8 4 9

2 1 4

5 3 0

3 1 7

2 1 6

50 60 0

30 20 54

20 20 46

8 11

2 0

6 11

0 0

4

4

5

6

0

50

0

45

50 55

5 10 21

0 2 5

5 8 16

2 1 3

2 7 13

1 2 5

40 10 14

40 70 62

20 20 24

27 12

9 4

18 8

5 0

14 5

8 7

19 0

52

30

42

58

14

4

10

2

7

5

14

50

36

8

1

7

1

3

4

13

38

50

10

0

10

1

7

2

10

70

20

20 16 5 21 22 15 2 0 280

0 1 4 2 0 2 0 0 69

20 15 1 19 22 13 2 0 211

6 8 2 9 10 5 0 0 65

6

8

30

5 1

50

40

30 19 40

40

3 2 2 3

10 9

43 45

10 14

2

8

33

13

2

0

0

100

48 41 53 0

0 113

0 102

0 0

0 0

0 0

31 20

Values in bold type mark the systems tract with the highest abundance for each MF type. MF types 9, 19, 23 and 24 are not considered, because of their minor occurrence. For the same reason, the distribution of MF types 2, 5, 7, 10 and 15 is not clearly significant. However, some distinct trends in their distribution are observable.

o f t e n rich in quartz, indicating t e r r i g e n o u s input into n e a r s h o r e areas. T h e macroscopically freq u e n t l y occurring dolomitic peritidal sediments are u n d e r - r e p r e s e n t e d by their n u m b e r of thin sections (Figs 4 a n d 5). (2) W a c k e s t o n e s and p a c k s t o n e s of p r o t e c t e d i n n e r r a m p to n o n - m a r i n e e n v i r o n m e n t s can be divided into (a) miliolid-rich w a c k e s t o n e s and p a c k s t o n e s ( M F type 17), which are widely dist r i b u t e d across the shallow shelf, excluding the n o r t h e r n m o s t sections, and (b) w a c k e s t o n e s and p a c k s t o n e s rich in ostracods ( M F type 18, Table 2), w h i c h are very c o m m o n especially in the

s o u t h e r n sections. Characteristic brackish-water o s t r a c o d associations are constituents of the L S T d e p o s i t s of s e c t i o n A N 2 (M. A. A. Bassiouni, pers. comm.). (3) Bioclastic p a c k s t o n e s (MF types 10-12) are of s u b o r d i n a t e i m p o r t a n c e within LSTs of all sections (Table 2). Oolitic bioclastic p a c k s t o n e s and grainstones of the shallow ramp ( M F types 2, 3 and 10) are less f r e q u e n t and limited to the n o r t h e r n area. It has to be c o n s i d e r e d that the n o r t h e r n sections (RN, Mi, R and M) do not r e a c h far into C e n o m a n i a n time. T h e i r facies types m a y thus be u n d e r - r e p r e s e n t e d .


MIDDLE CRETACEOUS SINAI RAMP

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Fig. 9. Simplified facies model and semiquantitative distribution of the different microfacies types (MFTs) across the shallow ramp transect during different systems tracts of Mid-Albian-Cenomanian times. (For classification and typification of the microfacies types, see Table 1.) The considered sections are indicated and arranged from N to S with respect to their lateral distances. The thickness of the beams indicates the abundance of the MF types. There are evident differences in abundance and distribution of the MF types within the different systems

Model. The lowstand carbonate ramp is mainly characterized by three facies belts (Fig. 10): (1) a proximal part emerged above sea level, where during M i d - A l b i a n - C e n o m a n i a n times non-

deposition or deposition of continental sediments prevailed; (2) a restricted environment also d e v e l o p e d in the proximal parts, where peritidal, lagoonal and brackish conditions


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occurred; (3) an open marine low-energy area, which separated the inner ramp from the midramp area (Fig. 10). The progradation of the facies belts and the emergence resulted in a strong decrease of the accommodation space. These lowstand conditions did not result in a simple shifting of the facies belts, but rather created a low-energy depositional system in the entire area of the shallow ramp, indicated by the obviously declining occurrence of grainstones. Furthermore, the stagnation of water circulation (compare Wright & Burchette 1996) resulted in the restricted depositional conditions observed and documented in wide areas during MidAlbian to Cenomanian time (Figs 9 and 10).

Transgressive systems tract The transition from LST to TST is marked by a total change of the hitherto common MF types. Now, the shallow ramp can be subdivided in two major facies belts that are mainly controlled by different hydraulic regimes. Within unit A, again represented only in the northern sections, bioclastic and oolitic grainstones (MF types 1, 2, 3 and 6) prevail together with bioclastic packstones (MF types 11,12 and 14), which are partly poorly winnowed (MF type 13) (Table 2). The grainstones are locally rich in terrigenous components (MF types 4 and 5) and in contrast to unit B, radial ooids occur. In general, all these environments are characterized

by MF types formed under high-energy conditions. In unit B, the distribution of MF types within the northern sections is similar to that for unit A. However, reworked rudist-bearing sediments (MF types 7 and 8) and oolitic as well as oncolithic packstones (MF types 10 and 11) are more common (Table 2, Fig. 9). In the southern parts, which were not covered by the sea during deposition of unit A, wide areas are now characterized by bioclastic wackestones to packstones with dasycladacean algae (MF type 16), typical of protected low-energy platform environments. Moreover, biopelpackstones and packstones to wackestones with benthic foraminifera (MF types 11, 12 and 14) characterize the southern regions. Subordinate rudist debris-bearing grainstones and rudstones occur in the north (MF type 7) and their micritic equivalents in the south (MF type 8).

Model. The TST depositional model shows significant differences from the LST model (Fig. 10). The main feature of TSTs are the large lateral extension of deposits representing carbonate shoals (northern region) and the predominance of open marine MF types in the proximal ramp areas (southern region). Thus, the distal inner ramp comprised a wide highenergy facies belt interfingering with an open marine low-energy area behind. The increasing accommodation space resulted in increasing

Fig. 10. Depositional models for the systems tracts of the Middle Albian-Cenomanian carbonate ramp. Most obvious differences concern the distribution of carbonate shoals and biostromes within the individual systems tracts.


MIDDLE CRETACEOUS SINAI RAMP water circulation and thus larger areas were affected by wave action, which allowed the development of extended oolitic and bioclastic shoals. Moreover, the efficient circulation had a positive feedback on carbonate production, which now became more effective in delivering sediment, when the sea level had risen enough to flood wide areas (Ginsburg in Hardie 1986). A similar relationship between increasing accommodation space and the formation of oolitic shoals was described by Jenkyns & Strasser (1995) and Pittet et al. (1995). Predominance of shoal grainstones in the TST is typical of ramps with high wave or tidal energy (Burchette & Wright 1992). The second main feature of the TST ramp in the northern Sinai is the large extension of an open marine low-energy facies belt (Fig. 10), which favours the occurrence of a dasycladacean-rich MF type. This cannot be explained by a simple shift of facies belts, because this MF type is never present within LSTs, although the shallow low-energy ramp is represented in several sections in the northern Sinai (Fig. 9). The increasing accommodation space and wave action during rising relative sea level additionally resulted in improved water circulation within the depositional area landward of the shoals, and thus open marine conditions prevailed in the southernmost region as well.

Highstand systems tract Again, the transition from TST to HST is characterized by some major changes in the distribution of MF types. Furthermore, the HST ramp shows a much better differentiation in facies belts than do the other systems tracts, resulting in an evident increase of MF type diversity. Within unit A, some grainstones with ooids and quartz or iron ooids still occur in the northern sections (MF types 3 and 4), whereas biopelmicrites were deposited in the central region (MF type 12). The sections (Figs 3 and 4) also show that dolomites deposited in peritidal areas are frequent and therefore under-represented by their number in thin sections. Unit B is characterized by a high amount of rudist biostromes in the north (Fig. 4). Grainstones containing ooids or bioclasts decrease drastically and are restricted to the northernmost region. Grainstones and packstones containing rudists debris (MF types 7 and 8), and floatstones with encrusters (MF type 9) constitute an important group of MF types (Table 2, Fig. 9). This high amount of rudist fragments coincides with the observed large extension of rudist biostromes within the outcrops. Bioclastic packstones locally with high amounts of

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pseudopeloids characterize the distal inner ramp (MF types 11-14), although the main abundance of these MF types does not lie within the HST (Table 2). Bioclastic wackestones and packstones with quartz, and dolomites with quartz are limited to the southern area (sections GM, H, A N 1 and A N 2), whereas the northern region was characterized by an environment free of siliciclastic input. The abundance of MF types 20, 21 and 22, which reflect tidal flat deposition, strongly increases in the late HST.

Model. For the HST, the Sinai ramp can be described by a low-energy depositional model (Fig. 10). The carbonate mud continuously delivered from the tidal flats and restricted inner ramp environments in the south was distributed all over the ramp areas, because the tidal flats were now enlarged owing to the beginning of progradation. Furthermore, the relatively lowenergy conditions prevented the export of carbonate mud, which together with the minor increase in accommodation space favoured the development of widely extended rudist biostromes. On carbonate platforms, the carbonate production generally reaches its maximum during HST, because of extended flooding and thus full carbonate production (Handford & Loucks 1993), and the formation of ooids is favoured during HST for the same reason (Schlager et al. 1994). On carbonate ramps, however, HST sediments are commonly grainier than those of the TST (e.g. Burchette et al. 1990) and commonly oolitic as well (Wright & Burchette 1996). In contrary, the Sinai ramp exhibits totally different conditions in the HST: owing to the prevailing progradational regime, the shoal complexes were shifted towards the north and resulted in topographical highs forming effective barriers to marine circulation (compare Handford & Loucks 1993). These barriers now caused environments with reduced wave and current activities and thus the end of effective ooid and shoal formation (Fig. 10). On the other hand, this HST scenario favoured the development of muddy inner ramp areas, and thus the formation of rudist biostromes, which acted as new barrier systems. In contrast to the LST, restricted conditions rarely developed on the low-energy HST ramp.

Discussion and conclusions Depositional model: climatic control The sedimentary features within systems tracts and bounding surfaces observed in the Upper Aptian to Cenomanian succession of the northern Sinai indicate that two sequence


276


btfatigiaphic models can be distinguished: the siliciclastical-influenced ramp for the Aptian to Lower Albian unit A, and the carbonate-dominated ramp for the Middle Albian to Cenomanian unit B. The differences in sequence stratigraphic features are most evident in deposits of sea-level lowstand, when the siliciclastic input may have been very high. We suggest that two main controlling factors causing the two different ramp settings have to be taken into account, as follows. Both a climatic change from humid to arid conditions as well as a general retreat of the coastline owing to a second-order sea-level rise would have resulted in the observed sequence stratigraphic patterns. During Aptian to Early Albian times, the prevailing deltaic sediments argue in favour of humid climatic conditions, confirmed by karstification (SB 8). We assume a less humid climate during Mid-Albian-Cenomanian times because of the disappearence of deltaic environments, the increased extension of tidal flats, and the less frequent karstification. Minor fresh-water influences, however, can still be recognized from the deposits of brackish environments in section AN 2. These observations are comparable with descriptions of 'tropical climates with semiarid background' recorded from the adjacent Western Desert during Aptian to Cenomanian times (Abdel-Kireem et al. 1996). Climatic changes have immense consequences for the sequence stratigraphic patterns, as demonstrated by Handford & Loucks (1993). However, although a less humid climate is envisaged for the upper part of the succession (Middle Albian to Cenomanian units), the change in climate was not drastic enough to be the only explanation for the changed facies patterns. The fact that really arid climatic conditions have not been proved from the succession (lack of evaporites and characteristic diagenetic features) and the assumption of tropical conditions in the adjacent areas (Western Desert, Abdel-Kireem et al. 1996) argue against an exclusive responsibility of climatic change. A general second-order transgression during Albian and Cenomanian time was demonstrated with respect to the southward retrograding coastline (Fig. 2; Kuss & Bachmann 1996). We assume a composite effect of these two factors to be the main reason for the changing sequence stratigraphic features. F o r m a t i o n o f ooids The high amounts of oolitic grainstones compared with bioclastic grainstones in the Sinai (Fig. 9, Table 2) is interpreted as the result of

increased development of flat, high-energy areas during the TST, which favoured the formation of ooids. Similar interrelationships between accommodation space and the formation of ooids have been mentioned by Jenkyns & Strasser (1995) and Pittet et al. (1995). Their observations of terminating ooid formation and synchronous increase of lagoonal and peritidal facies types coincide with our interpretations of TST to HST transitions, when decreasing accommodation space favoured the build up of carbonates close to sea level and the start of progradation. As a consequence, extended tidal flats and shallow lagoons occurred and coevally oolites were drastically decreasing (Fig. 10). The abundances and volumes of oolites within the studied succession are unusually high compared with other Aptian-Cenomanian periMediterranean platform settings of Cretaceous time (Jenkyns & Strasser 1995), such as some of the northern Tethyan margins, where no comparable oolites have been described (e.g. the Pyrenees: Lenoble & Canerot 1993; Vilas et al. 1993; Skelton et al. 1995; or the Apennines: Carannante et al. 1993). However, local oolitic grainstones have been reported from the Middle Cretaceous carbonate ramp of the Arabian Gulf (Alsharhan 1995), from southern France (Hunt & Tucker 1993), and from Arizona (Scott & Warzeski 1993). Although variations of accommodation space are obviously only one major factor controlling the formation of ooids, we would expect them in the platform environments of the northern Tethys as well, where higher rates of increasing accommodation space have been reported (e.g. Pyrenees: Lenoble & Canerot 1993); however, no evidence of ooids was proved. For the Sinai model we assume that the high amounts of oolitic grainstones (compared with bioclastic grainstones, Fig. 9) result from increasing accommodation space and the increasing extension of flat, high-energy areas during TST, which favour the formation of ooids in comparison with skeletal grains (Schlager et al. 1994).

S e q u e n c e stratigraphic patterns: controlling factors The major controlling factor for the formation of sedimentary cycles is the changing sea level, expressed by using the sequence stratigraphic concept and evidenced by emergence-submersion cycles reflected in emergence horizons and terrestrial sandstones alternating with marine sediments (Vail et al. 1991). As tectonic processes are negligible for the Sinai during Late


MIDDLE CRETACEOUS SINAI RAMP Aptian-Cenomanian times (Ayyad & Darwish, 1996), we assumed a constant rate of subsidence for the whole study area. Most sequence stratigraphic models for carbonate ramps assume that the relatively simple geometry of a homoclinal ramp basically means that facies belts simply move up and down the ramp in response to longterm relative sea-level changes (Tucker et al. 1993). Moreover, the sedimentation rates on carbonate shelves depend on the productivity in the marine-subtidal carbonate factory (Pratt & James 1986). The profile of total carbonate productivity against depth (Schlager 1981) clearly shows the influence of water depth, and consequently the width of the shallow-water zone is an important factor for carbonate production. The width of the sediment-production zone for a homoclinal ramp, however, is thus supposed not to change substantially with a fall or rise in relative sea level (Handford & Loucks 1993). Although this is true for many homoclinal ramps, the response of the Middle Cretaceous Sinai carbonate ramp to relative sea-level changes is different. The detailed facies analysis demonstrates that it is not possible to reconstruct one all-inclusive facies model; indeed, it is necessary to consider each systems tract separately (Fig. 10). Relative sea-level changes here did not result in a simple shift of facies belts; rather, they evidently affected the depositional environment with regard to siliciclastic input during Late Aptian-Early Albian time, and the current systems as well as biological productivity during Mid-Albian-Cenomanian time. Unit A . During Late Aptian-Early Albian times, the sporadic siliciclastic input during deposition of LSTs was detrimental for carbonate production and consequently authigenic sediments were produced primarily during TSTs and HSTs. Low growth rates of biostromes, deposits of restricted inner ramp environments and grainstones are common during the HSTs, as in platform environments (Schlager et al. 1994) and on many carbonate ramps (Elrick & Read 1991; Tucker et al. 1993). A similar model showing progradation of siliciclastic facies during lowstand of relative sea level, widespread carbonate deposition in keep-up settings during transgression of relative sea level, and high carbonate production resulting in progradation during HST has been described for Upper Devonian-Lower Carboniferous successions in the Canning Basin (Holmes & Christie-Blick 1993; Southgate et al. 1993). Comparable lithological changes from rudist-bearing limestones to continental sandstones or oyster-bearing marls have been attributed to regressions that were

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reported from a Late Albian-Mid-Cenomanian platform situated at the western margin of the Iberian Basin (Garcia et al. 1993). Unit B. The Mid-Albian to Cenomanian transition to a pure carbonate ramp system in the Sinai resulted in a depositional environment that responded to sea-level fluctuations with changing water circulation. Comparable variations in the circulation system are well known from platform environments, but they are not common on carbonate ramps (Handford & Loucks 1993). Here, the different palaeoenvironments, from the restricted LST ramp and the high-energy TST ramp to the low-energy open marine HST ramp, can be interpreted as a result of the changing water circulation system. The strongly falling sea level resulted in an exposure of the inner ramp or in restricted environments on the shallow ramp, both characterized by strongly reduced water circulation (Fig. 10). The formation of widespread, thick shoal sediments during TST can be interpreted as a long-term interval, where carbonate production lagged behind the sea-level rise or just reached the keep-up stage (compare Handford & Loucks 1993). Therefore, the accommodation space remained relatively large and favoured the creation of a high-energy milieu. A similar progressive deepening, indicated by transitions from lagoonal environments to bioclastic shoals, was described by Gimenez et al. (1993) from an Upper Albian-Cenomanian ramp at the Betic and Iberian Basins. To understand the development of prominent rudist biostromes during HSTs, we have to consider two factors. First, during constant relative sea level, carbonate production rapidly reached the catch-up stage and, from this time on, the steady progradation generated a low-energy, open marine ramp environment. Second, rudist associations were highly adapted to muddy, lowenergy environments. Thus, their growth maximum on the Middle Cretaceous northern Sinai ramp was reached during the HST.

The results summarized here have been obtained in the course of the DFG priority programme 'Global and regional controls on biogenic sedimentation' (Ku 742/10). Additional financial support was granted by the Peritethys Programme (No. 94-30). Technical assistance for the field work was given by members of the Ain Shams University, Cairo. We would like to express our special thanks to the following persons: M. A. A. Bassiouni (Cairo), M. Joachimski (Erlangen), J.-P. Masse (Marseille), M. Simmons (Aberdeen), M. Brinkmann (Bremen) and E. Friedel (Bremen). The manuscript benefited from suggestions given by P. Skelton and an anonymous reviewer.



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