International Journal of Earth Sciences

, Volume 105, Issue 5, pp 1387–1415 | Cite as

A new tectono-sedimentary model for Cretaceous mixed nonmarine–marine oil-prone Komombo Rift, South Egypt

Original Paper
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Abstract

The Komombo Basin is a recently discovered mixed nonmarine–marine, petroliferous basin of Cretaceous age in South Egypt. It is an asymmetrical half graben, synchronous with the Neothys opening and filled with up to 4 km of continental to open marine strata ranging from Early to Late Cretaceous. Despite its great relevance, no detailed sedimentological study concerning this basin has been carried out to date. Here, we present an integrated approach to the borehole and core data, as well as unique outcrop sections to construct a new detailed sedimentological interpretation on depositional systems, controls on basin evolution, basin configuration and regional tectonic setting. Seven depositional systems were recognized: (I) a fluvial fan system, (II) a braidplain system, (III) a siliciclastic lacustrine system, (IV) a lacustrine/lagoonal system, (V) a fluvial-estuarine system, (VI) a tidally affected delta, and (VII) an open marine system. The Komombo Basin evolution can be compartmentalized into three main rifting phases: the Berriasian–Early Barremian, Late Barremian, and Aptian–Albian. The first and third rifting phases are comparable with the rifting phases reported for several basins in North and Central Africa. The second rifting phase represents a transitional event between the other two phases. The first three depositional systems consist mainly of continental siliciclastics and are dominant in the Berriasian–Early Barremian and Late Barremian rifting phases. The lacustrine/lagoon and fluvial-estuarine systems correspond to the Aptian–Albian rifting phase, while the Campanian–Maastrichtian open-shelf deposits represents the post-rift stage.

Keywords

Cretaceous Nonmarine Rift Komombo Egypt 

Introduction and exploration history

The Komombo Basin is one of the recently discovered hydrocarbon-bearing basins, being located on the west bank of the Nile River, about 570 km southeast of Cairo, and about 260 km west of the Red Sea, Egypt (Fig. 1a). It is one of a series of intracontinental rift basins straddling the river Nile such as the Beni Suef and Asyut basins whose formation was proposed to be linked with breakup of Western Gondwana and the opening of the South and Equatorial Atlantic oceans during the Early Cretaceous (Bosworth et al. 2008). This caused active rifting episodes over large domains due to clockwise rotation of north and central Africa (Guiraud and Maurin 1992; Guiraud et al. 2005). Burke and Dewey (1974) proposed that Africa was subdivided during the Early Cretaceous into three main rigid blocks: the southern Austral Block, Arabian–Nubian Block, and Western Block (Fig. 1a). They refer to the boundaries between these blocks as major rift zones. The Early Cretaceous rifting is multiphase (Guiraud and Maurin 1992). The Late Berriasian to earliest Aptian (Maurin and Guiraud 1990; Guiraud and Maurin 1991, 1992; Guiraud et al. 2005) is characterized by ~E–W and NW–SE trending rifts initiated or developed over Central Africa and along the northern African–Arabian Tethyan margin. During the Early Aptian, a rapid change in the intraplate stress field occurred due to the migration of the Arabian–Nubian Block toward the northeast related to opening of the Indian Ocean (Guiraud and Maurin 1992; Guiraud and Bellion 1995). The extension direction, formerly N160 to N–S, moved to NE–SW (Guiraud and Maurin 1992). This resulted in rapid subsidence along the NW–SE trending rifts. In the Late Cretaceous, post-rift thermal sagging affected sedimentary basins both along the margins of the African plate as well as intracratonic basins (Bumby and Guiraud 2005). The rift geometry of Komombo Basin was detected from gravity data and outcrops (Kamel 1990). The Komombo Basin is filled with a Cretaceous mixed marine and nonmarine sedimentary succession which is over 4000 m thick in the deepest parts of the basin. However, many Mesozoic basins in South Egypt and Sudan contain dominantly nonmarine sediments (Nagati 1986; Schull 1988; Taha 1992). Early concepts considered the Komombo Basin formation is related to Jurassic rifting between Africa and Apulia (Nagati 1986; Kamel 1990; Dolson et al. 2001).
Fig. 1

a Tectonic setting of northeast Africa and eastern Mediterranean complied from Bosworth et al. (2008) and Guiraud et al. (2005). The paleoshorlines of Early and Late Cretaceous after Wycisk (1994). Indent map shows the reconstruction of West Gondwana breakup during the Early Cretaceous after Guiraud and Bellion (1995). A.B. Austral Block, A.N.B. Arabian–Nubian Block, W.B. Western Block. The white and black arrows refer to the motion of the Arabo–Nublian block during Early Barremian and Late Aptian, respectively (Guiraud and Maurin 1992). b Structural map of the study area showing NW–SE trending groups of faults and wells is also marked. The indent map shows structural trends of the Komombo Basin, SW Egypt (for location see a)

The study area is located on the central part of the Komombo Basin (Fig. 1b). It was part of a larger former Repsol concession which was relinquished in 2001 after the drilling of three wells in this area. These wells are located in a NW–SE trending half graben. After Centurion obtained exploration rights on Komombo Block in 2004, additional 2D seismic and 3D seismic acquisition and processing were done. Despite this, the basin represents unique petroleum activity in South Egypt, the geologic evolution, geometry, and petroleum system of the basin remain poorly understood.

The Komombo Basin was studied by Fathy et al. (2010), Wood et al. (2012), and Dolson et al. (2014). The aim of this study is integrated approach which describes and interprets the borehole image logs, dipmeter, wire-line logs and the available cores and combines them to establish depositional environments, paleosediment transport directions within these environments, and the tectono-stratigraphic development of the Cretaceous succession within a palyno-stratigraphically calibrated framework. The results were used as input for basin modeling and development purposes in this basin and related basins in North Africa; however, this aspect is not discussed here.

Data set and methodology

This study is based on results from the description and interpretation of core material from 6 wells, wire-line logs from 24 wells, and borehole images from 10 wells. The wire-line logs, cores, borehole images, and seismic data used in this study are courtesy of SeaDragon-Egypt. The detailed palynological description and biostratigraphic zonation were carried out for the studied wells by Abu El Ella (2006, 2011). The core and the borehole images were acquired from the same wells and are directly correlated. These data sets were first described and interpreted separately and subsequently integrated into one final interpretation. This interpretation led to the subdivision of the studied Cretaceous succession into facies based on: lithology, grain size, internal structures, color, and bioturbation. These facies are grouped into facies associations and depositional systems. Unique surface outcrops were used as analogs to the facies associations and depositional systems described from the subsurface data. These outcrops are located at the area between Faris Village and Kobet El Hawa (south of the study area, Fig. 1a). Paleogeographic maps were built using the combined primary suite of maps and incorporating interpreted depositional systems and facies associations. Each paleogeographic map reflects the distribution of depositional systems and facies at the time of maximum progradation. Maps also show the locations and nature of sedimentary elements, structural features and depocenters, axes of sediment input, and dominant transport. One seismic line from the 3D seismic survey has been used to show the basin architectures and lateral thickness variations of the sedimentary strata.

Geological setting

The Komombo Basin is a 70-km-long, 30-km-wide extensional rift basin of Early to Late Cretaceous age in southern Egypt (Fig. 1a, b). The Komombo Basin appears to be a half graben (Fig. 2a), probably developed in an extensional stress regime originated in Central and North Africa with the opening of South Atlantic during the Early Cretaceous (Bosworth et al. 2008). This extensional stress also caused reactivation for east-west Jurassic basins as in the Sarir Troughs of Libya and Faghur and Abu Gharadig basins of Egypt (Bosworth et al. 2008). The main bounding fault of the Komombo Basin is located to the northeast and trending NW–SE, while the subordinate faults are mainly synthetic and of the same trend (Fig. 1b). The basin contains up to 4 km of strata, principally clastics (sandstone and shales) with minor shelf carbonates in the upper part (Fig. 2b). The intracratonic rift basins form by the stretching and faulting of continental lithosphere (e.g., McKenzie 1978; Wernicke 1985; Lister et al. 1986; Rosendahl 1987; Kusznir et al. 1991; Ziegler 1992). Because rift basins form in a variety of settings, their sedimentary fill shows a great variability (Leeder 1995). Ideas on synrift sediment architecture have evolved from studies of nonmarine settings (e.g., Crossley 1984; Frostick and Reid 1987; Leeder and Gawthorpe 1987; Morley 1989; Lambiase 1990) and marine rift basins (e.g., Surlyk 1978, 1989; Leeder and Gawthorpe 1987; Prosser 1993; Gawthorpe et al. 1994). In addition to recognizing the asymmetric geometry of fault-controlled sedimentary packages, it has long been realized that changing lithological signatures and stacking patterns can reflect variation in fault-related subsidence rate, assuming a variation in the balance between sediment supply and subsidence rate. Clayprone intervals (with time-equivalent conglomerates restricted to narrow belts adjacent to the fault scarps) in syntectonic successions are known to represent periods of rapid differential subsidence (Steel 1988), whereas the intervals of coarse facies extending farthest from the fault scarp relate to periods of relative tectonic quiescence and minimal creation of accommodation space (Blair 1987). The Early to Mid-Cretaceous rifting phases were followed by tectonic Syrian Arc inversion in Late Cretaceous–Early Tertiary (Jenkins 1990; Sehim 1993; Moustafa and Khalil 1990; Ayyad and Darwish 1996; Guiraud and Bosworth 1997; Guiraud 1998; Guiraud et al. 2005; Bosworth et al. 2008). This tectonic inversion is associated with formation of folds and mid-basinal highs (Bosworth et al. 2008). During the Late Cretaceous–Tertiary, the Komombo Basin became epicontinental sea as a part of the southern Tethyan domain. During Late Neogene, the area was incised by the Nile Canyon with the formation of the thick Nile valley fill (Nile aquifer).
Fig. 2

a NE–SW seismic cross section through the study area showing the half-graben geometry of the Komombo Basin (Seismic data presented courtesy of SeaDragon-Egypt). b Correlation of the Cretaceous succession in the Komombo Basin. Northeast–southwest-oriented correlation of the deepest wells used in this study. Well spacing is to scale 1 = 70,000 (log data presented courtesy of SeaDragon-Egypt)

Stratigraphy

The stratigraphy of southwestern Egypt was discussed by Klitzsch and Squyres (1990) and Wycisk (1994). In South Egypt, the sedimentary succession of Carboniferous–Late Cretaceous comprises two main cycles: Karoo cycle (Carboniferous–Early Jurassic) and Nubian cycle (Late Jurassic–Cretaceous). The Karoo cycle was deposited in a shallow basin extending from South Egypt to northern Sudan and bordering areas of Chad and Libya (Klitzsch and Squyres 1990). The Western Desert of Egypt consists of a series of small rifts which were initiated in the Late Jurassic–Early Cretaceous (Nagati 1986; Guiraud 1998) such as Faghur and Abu Gharadig basins. The stratigraphy of the Cretaceous succession of the Komombo Basin (Figs. 2b, 3) includes 6 formations from base to top: Six Hills, Abu Ballas, undifferentiated Sabaya-Taref, Maghrabi-Quseir, and Dakhla formations. The Six Hills Formation is subdivided into seven members from base to top (A–G, Figs. 2b, 3): A member consists of gravelly sandstone with some mudstone interbeds. B member consists mainly of dark gray organic-rich mudstone. C–E members consist of successive alternations of pebbly to medium-grained sandstone and red to brown mudstone with general coarsening upward pattern. F member consists of variegated to red mudstone with thin sandstone interbeds. G member is similar to E member but shows wide spatial distribution. The Six Hills Formation represents Neocomian lowstand deposits and is characterized by noncyclic stable fluvial sandstone sheets (Wycisk 1994). The Six Hills Formation yielded well-preserved Imparadecispora apiverrucata, Cupressacites oxycedroides, and Aequitriradites spinulosus palynozone (Abu El Ella 2006, 2011, Fig. 3). Schrank and Mahmoud (1998) dated an Impardecispora/Aequitriradites assemblage as Berriasian–Valanginian. Aequitriradites spinulosus is restricted to Berriasian–Barremian at Siqeifa 1-X borehole, northern Egypt (Mahmoud and Deaf 2007). Impardecispora apiverrucata was found in the Saharan Neocomian subzone 5c (Reyre 1973). This zone is correlated with zones 1 and 2 of Schrank (1991, Fig. 3) and considered as Berriasian–Barremian (Abu El Ella 2006, 2011; Mahmoud and Deaf 2007). The species appeared later as Barremian–Aptian in South America (Argentina, Archangelsky and Gamerro 1967a, b). The younger Abu Ballas Formation (Aptian) represents highstand deposits and is composed of mudstones and siltstones of shallow marine settings (Wycisk 1994). It is characterized by appearance of angiosperm of Stellatopollis barghoornii and Stellatopollis spp. assemblage (Abu El Ella 2006, 2011). This assemblage was considered as Aptian–Barremian (Abu El Ella 2011; Mahmoud and Deaf 2007), while was considered as Barremian (Schrank 1992) and Albian–Aptian in Brazil (Regali 1989). Abu Ballas Formation unconformably underlies the Sabaya-Taref Formation (Albian/Early Cenomanian–Turonian). The Sabaya-Taref Formation is composed of cross-bedded to rippled sandstone with thin mudstone interbeds of fluvial to marginal marine setting. Some variegated mudstones occur at the transition from the Sabaya fluvial sandstone to the tidal-dominated Taref sandstones. These mudstones are considered as Maghrabi Formation. The Albian Sabaya Formation is characterized by cyclic nature and increase the soil formation by the sea level rise, while the Early Turonian Taref Formation represents lowstand stable fluvial sheets (Wycisk 1994). The Sabaya-Taref Formation is yielding palynomorphs such as Classopollis spp., Inaperturopollenites sp., and Crybelosporites pannuceus with miospore assemblage is represented by Baculatisporites sp., and Spheripollenites sp. (Abu El Ella 2011). Crybelosporites pannuceus represents Aptian palynological records in Sudan (Kaska 1989). Younger records of Crybelosporites pannuceus in Egypt, as Albian–Cenomanian, are documented by Schrank and Mahmoud (1998), Mahmoud and Deaf (2007), and Abu El Ella (2011). It appeared also later (Albian) in other African localities (NE Nigeria, Lawal and Moullade 1986; NE Libya, Batten and Uwins 1985), in South America (Brenner 1968), and in the Middle East (Qatar, El Beialy and Al Hitmi 1994). The overlying Turonian–Santonian Quseir Formation consists mainly of light gray shales with thin sandstone intercalations. The contact between the Quseir Formation and the underlying Sabaya-Taref formations is gradational. The Quseir Formation is characterized by upward increase in the sandstone interbeds toward the top. The upper part is dominated by dark gray mudstone with some coal streaks (Coal marker). This formation is characterized by the dominance of Foveotricolpites gigantoreticulatus, Foveotricolpites giganteus, and Droseridites senonicus in most of the investigated samples as well as occurrence of Ephedripites <30 μm and Equisetosporites ambiguous (Abu El Ella 2011). This zone can also be correlated with zone 7 (Schrank 1992) and zone S–V (Awad 1994) from Egypt and the Sudan, both dated as Turonian–Santonian. This zone can be compared with the middle assemblage of Mohsen (1992) described from the Kharga Oasis, dated as Coniacian, zone VII (Schrank and Ibrahim 1995) and zone 5 (Ibrahim 1996), both dated as early Turonian from the north Western Desert. Ibrahim and Abdel-Kireem (1997) dated similar zone as Turonian–Santonian from Farafra Oasis. The Campanian–Maastrichtian Dakhla Formation occupies a stratigraphic position at the top of the Quseir Formation. It consists of dark gray to black shale, marl, and clay with intercalations of calcareous, sandy, and silty beds (Said 1990). It is characterized by dominance of dinocyst such as Andalusiella polymorpha, Paleocystodinium australinum, Paleocystodinium gabonense, Senegalinium bicavatum, Batiacasphaera reticulata, Chytroeisphaeridia baetica, Deflandrea obliquus, Trichodinium castanea, Senegalinium granulostriatum, and Trichodinium sp. (Abu El Ella 2006, 2011). This zone is correlated with zone 8 of (Schrank 1992). Senegalinium bicavatum is the most characteristic species of the Campanian–Maastrichtian of Senegal, Gabon, Morocco, Brazil, and Egypt (Jain and Millepied 1973; Boltenhagen 1977; Malloy 1972; Herngreen 1975; Rauscher and Doubinger 1982; Schrank 1987). The Dakhla Formation is incised by the Nile Canyon with dominance of sands and gravels (Nile Fill, Figs. 2b, 3).
Fig. 3

Litho-, bio-, and tectono-stratigraphic framework of the study area. Comparison of the palynozones of Abu El Ella (2006, 2011) with previous schemes in NE Africa (Schrank 1991, 1992) and N Morocco (Gübeli et al. 1984)

Lithofacies and depositional environments

The interpretations presented here for the Cretaceous Komombo Basin filling succession are based on detailed sedimentological analysis, detailed stratigraphic correlation, lithofacies, and paleocurrent analyses. The results suggest that the whole succession was deposited in seven depositional systems (Fig. 2b), some of them contemporaneous and some succeeding each other during basin evolution. These are: (I) a fluvial fan system, (II) a braidplain system, (III) a siliciclastic lacustrine system, (IV) a lacustrine/lagoonal system, (V) a fluvial-estuarine system, (VI) a tidally affected delta, and (VII) an open marine system. The fluvial fan system comprises two facies associations: (Ia) stream-dominated fan and (Ib) floodout-dominated fan. The braidplain system can be described in terms of three facies associations: (IIa) distal, (IIb) mid-, and (IIc) proximal braidplain. The lake/lagoonal system, siliciclastic lacustrine, and open marine systems, each one encompasses one facies association. The fluvial-estuarine system includes five facies associations: (Va) mixed-load fluvial channels, (Vb) tidally affected fluvial channels, (Vc) interfluve, (Vd) tidal channels, and (Ve) tidal-flat deposits. The tidally affected delta system includes two main associations: (VIa) a deltaic mouthbar and (VIb) deltaplain deposits.

The facies associations (FA) are related to the first three depositional systems, essentially constituted of continental siliciclastics, forming the Six Hills Formation. The lake/lagoonal depositional system composes the Abu Ballas Formation, while fluvial-estuarine system composes of undifferentiated Sabaya-Taref formations and delta depositional system composes the Quseir Formation. The open marine system forms the Dakhla Formation. The facies association (FA) was built up of one or more of sixteen facies (F) whose details are given in Table 1.
Table 1

Description and interpretation of the lithofacies

Facies

Facies description

Interpretation

F1—Massive conglomerates and conglomeratic sandstone

Amalgamated and stacked discontinuous ribbons and continuous sheets, 10–50 cm thick, internally organized as fining-upward pattern with many internal reactivation surfaces (Fig. 6a)

Product of destabilization of upper-flow regime antidunes during sheet flood on alluvial fans

F2—Massive pebbly muddy sandstone to muddy pebbly sandstone

Decimeter to meter thick bodies of pebble to mud grade, poorly sorted, structureless sandstones

Rapid deposition and/or destruction of structure due to bioturbation (c.f. Cain and Mountney 2009)

F3—Massive sandstone

It is made of decimeter to meter scale thick of medium- to coarse-grained sandstone to pebbly sandstone. It rests on basal scour

Deposition by direct fallout sedimentation from heavily sediment-laden flows during waning floods (Todd 1989; Maizels 1993; Jo et al. 1997)

F4—Stratified sandstone

Medium- to coarse-grained, sandstone to pebbly sandstone locally conglomerates with plane-parallel and low-angle cross-beds composing apparently tabular meter-scale successions often associated with F13

Migration of straight-crested dunes

F5—Cross-stratified sandstone to pebbly sandstone

Medium- to coarse-grained, small to medium-scale trough cross-stratified sandstones with sparse granules and pebbles disposed in erosive-based 10–200-cm-thick lenses, internally organized as 5- to 25-cm-thick sets

Channelized lower-flow regime deposits formed during recessional stages of sheet-flood events migration of sinuous-crested dunes

F6—Inverse-graded sandstone

Inverse-graded tabular beds of very coarse sandstone,

locally conglomerate, to fine-grained sandstones with a few plane-parallel laminations (Fig. 8c). Beds compose centimeter to meter scale

Rapid dumping from flood-dominated river discharge

F7—Wavy-bedded sandstone

Decimeter to meters-scale thick, light gray to brownish gray medium- to fine-grained cross-stratified sandstone with mud drapes along the cosets and reactivation surfaces. It ranges from flaser to lenticular bedding in nature

Cross-stratified sands are the product of dune/ripple migration (Ashley 1990) in the subtidal part of tidal-fluvial channels (Choi et al. 2004)

F8—Muddy sand heterolithics

Repetitive alternations of fine sand and shale laminae The fine sand and silt laminae range from millimeter-thick lenses (‘linsens’ of Reineck and Wunderlich 1968), to thicker rippled beds (less than a centimeter thick) showing planar and trough cross-bedding

Tidal rhythmites (Smith et al. 1991)

F9—Inclined muddy sand (HIS)

Sand layers are composed of very fine sand, varying in thickness from few millimeters to up to 30 cm and show upper-flow regime parallel lamination or current ripple lamination with opposite directions of migration

The IHS is particularly abundant in tide-influenced settings (Smith 1987, 1988; Thomas et al. 1987; Eberth 1996; Dalrymple et al. 2003; Choi et al. 2004) as a result of lateral meander migration. Point bars may develop IHS in the inner bends of meandering channels

F10—Contorted mudstone and sandstone

Mudstone and fine- to medium-grained sandstone beds with convolute folds (Fig. 9a)

Product of slumping

F11—Massive and laminated red mudstones

Decimeter to meter scale thick, of red to reddish brown mudstone that can be either massive or laminated (Fig. 8b). The laminated variety usually presents sparse sand grains and some sandstone interbeds

Deposition from suspension fallout under quite oxidized conditions

F12—Variegated sandy mudstone

Gray to light gray sandy mudstone to muddy sandstone with remains of plant debris

Suspension fallout in quite conditions as channel abandonment or lake

F13—Gray organic-rich mudstones

Decimeter to meter scale thick, of gray to dark gray, organic-rich mudstone that can be either massive or laminated (Fig. 8b). The laminated variety usually presents sparse sand grains and some sandstone interbeds

Deposition from suspension fallout in quite reducing water column

F14—Thinly laminated mudstone and siltstone

Dark gray to gray laminated to bioturbated mudstone and siltstone with some traces as planolites, teichichnus, and skoliths

Suspension fallout from oscillating tidal currents in flooded bays or tidal flats

F15—Laminated shale

It is made of thinly laminated shales and siltstone with dominance of marine dinocycts

Suspension fallout in open marine conditions.

F16—Coal to coaly mudstone

Coals and carbonaceous shales are easily identified on the open-hole logs because they show distinctly different log signatures from the sandstones and mudstones

Accumulation and preservation of plant material over a prolonged period in a swamp on a poorly drained floodplain

Fluvial fan system I

Fluvial fan deposits occur in the central part of the Komombo Basin. They comprise two main facies associations: (Ia) stream-dominated FA and (Ib) floodout-dominated FA. The facies association (Ia) dominates the axial depocentral occurrences, while the facies association (Ib) dominates the transverse hanging wall ones. These deposits are characterized by coarse-grained facies, mainly conglomerates and conglomeratic sandstones, and are characteristic of the rift initiation phases.

Stream-dominated mid-fluvial fan FA (Ia)

This FA is well developed in depocentre of Komombo Basin and nearby the master fault of Komombo Basin. It is dominated by successive fining-upward cycles of erosively based, massive, cross- to horizontally stratified conglomerates to conglomeratic sandstone facies (F1, F3, F4, and F5) with thin light gray to variegated sandy siltstone and sandy mudstone (F12, Fig. 4a, b). Petrographically, the conglomeratic sandstone consists of quartz grains being cemented with iron oxy-hydroxide (Fig. 4c). The F12 facies at top of the cycle is characterized by scattered quartz grains in siliciclastic clay and silt matrix (Fig. 4d).
Fig. 4

Examples of the stream-dominated mid-fluvial fan facies association with member A of the Six Hills Formation at well BHK-2. a Facies stacking and dip pattern and petrophysical log response; gamma ray (GR) and density/neutron (RhOB/NPHI). b Core photographs and sedimentologic log of the six-foot cored interval between 8462ftMD and 8468ft. c Microscopic view of the cross-startified pebbly sandstone (F5), plane polarized light. d Microscopic view of the variegated mudstone facies (F12), plane polarized light (core and microscopic photographs presented courtesy of SeaDragon-Egypt)

The common occurrence of scour-based, meter-scale fining-upward cycles of facies F1, F3, F4, and F12 reveals that the main process responsible for the deposition of this facies association was characterized by large variations in discharge. These cycles are interpreted as deposits of high-energy floods, which started with the erosion of previously deposited sediment at peak discharge and generated different sedimentary facies at different stages of the waning flow. The basal, channel-form scours are interpreted as the product of lower-flow-regime currents developed in the deeper portions of the stream. Above the scour fill, deposition was controlled by the migration of three-dimensional dunes (F5) and local gravel bars (F1), both in the lower-flow regime, suggesting that deposition took place inside confined streams. This was followed by waning of the floods and deposition of fine-grained facies (F12). The ephemeral nature of the streams is suggested by the recurrence of these cycles, marked by episodic peak discharge followed by waning flow, with no evidence of long-lasting active channels. This association is interpreted as fluvial channel deposits intercalated with overbank muds or playa muds probably of fluvial fan setting. Similar association was described by Ito et al. (2006) and Almeida et al. (2009).

Floodouts-dominated distal fluvial fan FA (Ib)

Floodouts-dominated fluvial fan deposits are well developed in the western part of the Komombo Concession as lateral drainage system at wells BHK-19 and BHK-20 (Fig. 1). Facies association Ib consists of successive alternations of massive to crudely stratified pebbly muddy sandstone to muddy pebbly sandstone (F2) and gray, massive mudstones (F12, Fig. 5a, b). Petrographically, the F2 shows variable matrix composition from sand-rich matrix with little clay content (Fig. 5c) to high clay content with little sand matrix (Fig. 5d). This facies association shows coarsening and thickening upward cycles of decimeters to meters scale.
Fig. 5

Examples of the floodout-dominated distal fluvial fan deposits of member A of the Six Hills Formation at BHk-19 well. a Core photographs of the muddy gravelly sandstone facies. b Close-up view for the core slab of the muddy gravely sandstone (F1). c Microscopic view of a muddy gravelly sandstone (F1) sample, plane polarized light (core and microscopic photographs presented courtesy of SeaDragon-Egypt)

The dominance of massive to weakly laminated fabric and poorly sorting fabric of F2 indicates rapid deposition or destruction of structure due to bioturbation (c.f. Cain and Mountney 2009). This facies is attributed to the diffuse flow of unchannellized flood waters across the alluvial plain that probably emanated either from breaches of the main channel bank or from points where fluvial channels terminated on the alluvial plain (cf. Tooth 2000a, b, 2005). This facies is probably interpreted as frontal lobe or floodout deposits fed by the terminal parts of channels or alluvial fan as the flow expanded out onto the floodbasin. The flood deposition was culminated by dominance of floodbasin to lacustrine deposition from suspension fallout as seen from dominance of the muds. The general successive coarsening and thickening upward indicates general progradation of the floodout probably at the distal part of alluvial fan toward the depocenter of the basin with the rotation of the hanging wall (c.f. Leeder and Gawthorpe 1987).

Braidplain depositional system II

The braidplain depositional system comprises three main facies associations: (IIa) distal braidplain, (IIb) mid-braidplain, and (IIc) proximal braidplain. These three facies associations are defined, with contrasting sand/shale ratios and similar sedimentary structures.

The distal braidplain FA (IIa) is dominated by red, massive to laminated mudstone (F11) with some successive encased lenticular bodies and sheets of cross-stratified and horizontally stratified medium- to coarse-grained sandstones (F4 and F5, Fig. 6a). These sandstone facies are arranged in erosively based fining-upward pattern (Fig. 6b). The dominance of mudstones reflects the deposition from suspension fallout under low-energy and oxidizing conditions in distal braidplain. Red coloration of sediments requires oxidizing conditions, preferably in a tropical to subtropical climate with a distinct dry season (Duchaufour 1982; Kampf and Schwertmann 1982; Birkeland 1984). The cyclic fluid flow interrupted the quite conditions with formation of ribbon channels.
Fig. 6

Examples of the distal braidplain facies association with the C member of the Six Hills Formation at well BHK-4. a Facies stacking and dip patterns (in dip track and stereographic projection), and petrophysical log response (GR, RhOB/NPHI). Close-up view of the BHI log of the cross- and horizontally stratified sandstone (F4 and F5) and red mudstone alternations (F11). Notice the southerly paleocurrent direction and scour at the sand base reflect the erosive action (channels) (log data presented courtesy of SeaDragon-Egypt)

The mid-braidplain FA (IIb) consists of semiequal alternations of cross-stratified and horizontally stratified coarse-grained to pebbly sandstones (F4 and F5) and red, massive to laminated mudstone (F11). It represents transitional facies between the proximal and distal braidplain deposits (Fig. 7a, b) where the sandstone bodies become thicker and extensive laterally (Fig. 7a, b).
Fig. 7

Examples of the mid-braidplain facies association with the C member of the Six Hills Formation at BHK-4 well. a Facies stacking, dip patterns (in dip track and stereographic projection), and petrophysical log response (GR, RhOB/NPHI). b Close-up view of the BHI log of the cross- and horizontally stratified sandstone (F4 and F5). Notice the southerly paleocurrent direction and scour at the sand base reflects the erosive action (possible channels) (log and image data presented courtesy of SeaDragon-Egypt)

The proximal braidplain FA (IIc) consists of erosively based sandstones, conglomeratic sandstones, and conglomerates (F1–F5) with subordinate mudstone (F11, Fig. 8a). The clasts consist of quartz and pink granites. Conglomerates, having grain sizes ranging from 1 to 5 cm, have erosive bases and are interpreted as channel lag deposits. This facies association shows wide distribution in the study area in wells and crop out in the study area to the south at Kobbet El Hawa (Fig. 8b). This facies association represents deposits of braided channel belt. The channel belt deposits are characterized by the spatial association with cross- and horizontally stratified and massive sandstones and conglomerates in broad lenses often organized in decimeter to meter-scale fining-upward cycles with erosional bases (Fig. 8a–c). The erosional contacts of the cycles are the result of peak discharge currents acting upon previously deposited sediment in deeper portions of the stream channels. Deposition of the bedload due to loss of transport capacity during the waning stages of floods resulted in conglomerate bed sets concordant with the underlying erosional scours. Above the scour fill, deposition was controlled by the migration of gravel and sand bedforms, in subcritical flow streams, with possible development of upper-flow regime plane-parallel bed due to water depth reduction during the waning flow. The cross-stratified sandstone reflects the 3D and 2D dunes migration at shallow broad channels. The mudstone facies is interpreted as floodplain muds or shallow lacustrine deposits.
Fig. 8

Examples of the proximal braidplain deposits of the D member of the Six Hills Formation, well BHk-19. a Facies stacking and dip pattern and petrophysical log response (GR, RhOB/NPHI). b Surface analog for the proximal braidplain facies deposits at south west of the study area. Notice the stacked and amalgamated sand ribbons with many internal scour surfaces (log data presented courtesy of SeaDragon-Egypt). c Screen snapshot of the BHI log display in the interactive dip and facies picking screen of techlog (Schlumberger). The snapshot includes from left to right; GR (gamma ray), the left-hand image is static normalized (less contrast), right-hand BHI image is dynamic normalized and the last is the manually picked cross bedding. To the right stereographic summary of all measured cross-stratified and low-angle laminated sandstones in the entire interval. (Log data presented courtesy of SeaDragon-Egypt)

Lacustrine depositional system III

Lacustrine deposits are generally well developed in the area nearby the depocenter beside the master normal fault. It consists mainly of gray and dark gray partly laminated, organic-rich mudstones (F13, Fig. 9a, b). This kind of facies is very distinctive in the context of the B member of the Six Hills Formation. The fossil content lacks the marine fossils and is dominated by land-derived miospores such as Imparadecispora apiverrucata and Cupressacites oxycedroides, Callialasporites trilobatus, Inaperturopollenites sp., Spheripollenites sp., Classopollis sp., Cyathidites minor, Gleicheniidites senonicus, Cyathidites australis, Concavissimisporites sp., Cycadopites sp., Gleicheniidites sp., and Gleicheniidites delcourtii (Abu El Ella 2011). The lack of the marine fossils and the fine-grained nature of this association as well as dominance of land-derived miospores suggest deposition from suspension in stagnant water, below the depth of wave reworking of sand, in a lacustrine environment. Grayish black shales and dark mudstones reflect anoxic and reducing water conditions. The B member of Six Hills Formation is of Hauterivian–Early Barremian age, based on palynological studies by Abu El Ella (2011).
Fig. 9

Examples of the siliciclastic lacustrine facies of the B member of the Six Hills Formation at well BHK-19. a Core photograph of the gray organic-rich mudstone facies (F13). b Close-up view of F13. Notice the cemented fractures (arrows) (core photographs presented courtesy of SeaDragon-Egypt)

Lacustrine/lagoonal system IV

The lacustrine/lagoonal system represents the stratigraphic boundary between the nonmarine and marine deposits (Aptian) of the Komombo Basin. It has well-preserved thickness at the basin depocenter at well BHK-1 (Fig. 2b). It thins out toward the southwest and northwest. The deposits of this system consist of variegated to gray laminated and massive mudstones and siltstone facies (F14) and subordinated fine-grained sandstone especially toward the basin margins (Fig. 10a, b). The nature of this waterbody is not clear, since there is no sedimentological evidence favoring either a lake or lagoon. The Palynoflora from this facies association are dominated by land-driven pollens and spores as Stellatopollis sp., Foveosporites sp., Myssapollenites sp., Laevigatosporites sp., Leptolepidites sp., Todisporites sp., Staplinisporites sp., Sterisporites sp., and very rare Dinoflagellates as Odontochitina proifera (Abu El Ella 2011). Previous works suggest a marine environment based on global correlations, but a lacustrine environment is considered a strong alternative.
Fig. 10

Examples of the lacustrine/lagoonal deposits of the Abu Ballas Formation (upper part) and proximal to mid-braidplain of G member, Six Hills Formation (lower part) at well BHk-4. a Facies stacking and dip pattern and petrophysical log response (GR, RhOB/NPHI). Notice the northerly dipping of sandstone stratification. b Close-up view for the dynamic normalized BHI log shows the rhythmic alternations of massive and laminated mudstone and siltstone (log and image data presented courtesy of SeaDragon-Egypt)

Fluvial-estuarine depositional system (V)

The fluvial-estuarine depositional system is widely distributed in the Komombo Basin through the Upper Cretaceous, and it is penetrated in all the studied wells. It includes a wide spectrum of the mixed-load fluvial channels (Va), tidally affected fluvial channels (Vb), interfluve (Vc), tidal channels (Vd), and tidal-flat (Ve) facies associations.

Mixed-load fluvial channels FA (Va)

Mixed-load fluvial channels FA consists of cyclic fining-up (Fig. 11a), each one begins with erosively based, cross-stratified, fine- to medium-grained sandstone (F5, Fig. 11b, c). Thinly interbedded siltstones and mudstones are common through these deposits (F12, Fig. 11a). This association is interpreted as having been deposited in fluvial channels. A channel-fill interpretation is supported by the presence of erosional basal contacts, fining-upward trends, and upward decrease in bedform size. The presence of planar cross-stratification indicates migration of unidirectional, subaqueous dunes. The sandstone facies represents fluvial channel fill of relatively large and deep, low to high sinuosity channels. The abundance of facies F12 suggests a calm-water environment as overbank deposits, where settling of fine-grained particles took place. The rhythmic intercalation of F12 with the sandstone facies (F5) indicates cycles of sand input, both from traction and suspension under subaqueous conditions then probably subaerial exposure. The presence of the reactivation surfaces reflects the successive flows rather than single flood event. These fluvial channel sandstones are laterally continuous and are associated with well-developed overbanks.
Fig. 11

Examples of the mixed-load stream facies deposits of the Sabaya-Taref Formations, well BHk-19. a Facies stacking and dip pattern and petrophysical log response (GR, RhOB/NPHI) of mixed-load stream deposits of the Sabaya Formation. b Close-up view for the BHI log include the left-hand dynamic normalized and right-hand manually picked sedimentary structures. c Core photographs of the cross-stratified sandstone facies. d Core photograph for the slumped cross-stratified sandstone with mud drapes (log data and core photographs presented courtesy of SeaDragon-Egypt)

Tidally affected fluvial channels FA (Vb)

This facies association consists of fining-up, light gray to yellowish gray, fine- to medium-grained, calcite-cemented, cross-stratified and wavy-bedded, partly massive sandstones (F7, F3, and F5) and some gray mudstone interbeds (F14) at the top (Fig. 11a, b). Slumped or contorted cross-stratified sandstone (F10) is described at the base. Some foresets are delineated by small, flattened coal intraclasts. Stylolites and rare, very thin mudstone drapes occur toward the top (Fig. 11d). The scale of the cross-bed sets decreases toward the upper part, and the bi-directional cross-stratification can be observed (Fig. 11b). The presence of planar cross-stratification indicates migration of unidirectional, subaqueous dunes. Fining-up trends and heterolithic facies reflect gradual decrease in current speed from deeper parts of the channel to the channel bank (Mossop and Flach 1983). Abundance of mud drapes suggests slack-water periods during high- or low-tide stillstands. This together with opposingly directed ripples and herringbone cross-stratification reflects a strong degree of tidal influence, typical of the fluvial-tidal transition (Dalrymple and Choi 2007; Van den Berg et al. 2007). The diagnostic features of this facies are diverse soft-sediment deformation structures (slumps, ball-and-pillow, pseudonodules, and fluid-escape features).

Interfluve areas FA (Vc)

Interfluve areas FA (Vc) consists mainly of gray to variegated massive mudstone (F12) with rootlets and biogenic criteria (Fig. 12a, b). It is barren of dinocysts and palynoflora. This facies is interpreted as paleosol, which most likely developed in interfluve area. The rootlets and the red coloration reflect subareial exposure.
Fig. 12

Examples of the estuarine deposits of the Sabaya-Taref Formations, well BHK-19. a Facies stacking and dip pattern and petrophysical log response (GR, RhOB/NPHI) of tidal-dominated deposits of the Taref Fm. b Core photograph for the tidal channel deposits. c Core photograph of the successive coarsening upward tidally affected bars (yellow arrows). d Core photograph of the cross-laminated sandstone facies (F5) and inclined heterolithics (F9) of the tidal channel deposits (log data and core photographs presented courtesy of SeaDragon-Egypt)

Tidal channels FA (Vd)

This association includes tidal channels and tidal bars. The tidal bar category is characterized by successive coarsening upward pattern (Fig. 13a–c), while the channel deposits show fining-up pattern (Fig. 13a, d).
Fig. 13

Examples of the interfluve paleosol deposits of the Maghrabi Formation, well BHK-9. a Core photograph for the interfluve paleosol. b Close-up view for the massive variegated mudstone (F12) of the interfluve paleosol (core photographs presented courtesy of SeaDragon-Egypt)

The constituents of this association are mainly cross-stratified, fine- to medium-grained and wavy-bedded sandstones with mudstone intercalations (F3, F7, F8, and F14, Fig. 13b–d). Sandstones up to 15.2 m thick are cross-stratified, contain single and double carbonaceous drapes on foresets, and are slightly bioturbated (Fig. 13). Palynoflora from mudstone interbeds in the thick sandstone is dominated by land derived species as Stellatopollis barghoornii and Stellatopollis spp., and rare dinoflagellates (Abu El Ella 2011). One of the most common facies recorded in this association is the inclined heterolithics (F8). It consists of light gray, very fine- to fine-grained sandstones interbedded with dark gray to black siltstones (Fig. 13b). Wavy bedding and flaser bedding (single, double, and bifurcated) are dominant, in addition to lenticular bedding (with connected or single, thick lenses) being very common (Fig. 13b–d).

The fining-upward nature and the presence of the basal scour reflect the channel regime. Planar cross-stratification probably indicates ripple or dune migration across the channel floor. The occurrence of the heterolithics and mud drapes indicates a tidal influence on the depositional environment. Mud drapes record countercurrent flows and slack-water conditions, or weak tidal inundation (Fischbein et al. 2009). Double mud drapes and bundle thickness variations in cross-bed sets evidence the in-shore tidal environment of deposition (Visser 1980; Nio and Yang 1991). The inclined-heterolithic stratification is interpreted as having been produced by point-bar accretion, a dominant structure in upper-intertidal channels (e.g., Reineck 1958; Bridges and Leeder 1976; De Mowbray 1983; Thomas et al. 1987).

Tidal-flat FA (Ve)

Tidal-flat FA consists mainly of gray, laminated, and bioturbated mudstones with some thin, lenticular, and wavy bedded, rippled very fine- to fine-grained sandstones (F7, F8 and F14, Fig. 14a, b). This facies is interpreted as having been deposited in a restricted tidal-flat environment, most likely representing the lower-intertidal sand flat developed in a lower- to middle-estuarine setting. Presence of cross-lamination dipping in opposite directions suggests most probably to tidal influence with flow reversals. Bedload transport during tidal flow and suspension settlement during slack-water periods are indicated by alternating flaser and wavy bedding (Reineck and Wunderlich 1968; Klein 1971). Local presence of planar cross-lamination is ripple migration during periods of high-energy currents (Dalrymple 1992). Open marine, tidal-flat deposits commonly exhibit abundant and diverse biogenic structures (e.g., Mángano et al. 1996a, b). The paucity of body and trace fossils in this association supports the interpretation of a tidal flat developed in a restricted embayment. Absence of root traces and pedogenic slickensides suggests subaqueous conditions.
Fig. 14

Examples of the tidal-flat deposits of the Quseir Formation, well BHK-9. a Facies stacking and dip pattern and petrophysical log response (GR, RhOB/NPHI) of tidal-flat deposits. b Core photograph for the tidal-flat deposits shows the sandy mudstone heterolithcs. Notice the dominance of trace fossils as planolites (pl) and teichinus (te) (log data and core photographs presented courtesy of SeaDragon-Egypt)

Tidally affected deltas system (VI)

Tidally affected deltas are widely distributed in the Komombo Basin in the subsurface and the outcrop sections. It is well recorded at the top of the Coniacian–Santonian Quseir Formation. It includes two main facies associations: prograding mouthbar and deltaplain deposits.

Prograding mouth bar FA (VIa)

The mouth bar facies association comprises planar cross-stratified and inverse-graded, fine- to very fine-grained sandstones (F6) and laminated mudstone intercalations (F15) with general coarsening and thickening upward (Fig. 15a, b). This association is characterized by abundance of land-derived miospores relative to dinocysts (Abu El Ella 2011). The general dip data reflect the northward paleocurrent direction. This facies is interpreted as deltaic mouthbar as confirmed by the clear increase in the land-derived miospores relative to the dinocysts.
Fig. 15

Examples of the sandy mouthbar deposits of the Quesir Formation, well BHK-5. a BHI image log includes from left to right depth track, static BHI image, dipmeter track, and dynamic BHI (more contrast). b Surface analog for the sandy mouth bar at south west of the study area. Notice the successive coarsening upward cycles of the mouthbar deposits (yellow arrows). The author is the person for scale (log and image data presented courtesy of SeaDragon-Egypt)

Delta plain FA (VIb)

It is composed of two main subassociations: deltaplain muds and deltaic channel. The deltaplain muds subassociation consists of carbonaceous mudstone, coal, and siltstone intercalations (F14 and F16). Coal and carbonaceous shales are easily identified on the open-hole logs because they show distinctly different log signatures from the sandstones and mudstones (Fig. 16a). The coals are characterized by the lowest RHOB (density log) and highest NPHI (neutron log) in the study wells in addition to resistive appearance on the BHI log (Fig. 15b). According to Fielding (1987), three principal conditions are required to form peaty organic residues, namely (1) sufficient vegetation growth, (2) reduced oxidation and suppression of bacterial reduction probably by an elevated water table, and (3) reduced clastic input. With these conditions in mind, the absence of rootlets in this facies subassociation may be indicative of peat sourced from mires, or alternatively introduction of plant material by fluvial processes.
Fig. 16

Examples of the deltaplain deposits of the Quseir Formation (image data presented courtesy of SeaDragon-Egypt)

The deltaic channel subassociation is composed of cross-stratified, fine- to medium-grained sandstone with general fining-upward pattern (Fig. 16a). It rests on erosional base (Fig. 16b), and it may erode the underlying mouth bar deposits.

Open marine system (VII)

The open marine system consists mainly of rhythmic alternations of gray to dark gray, laminated shale and siltstone (F15) with thin sandstone interbeds (Fig. 17a, b). It is characterized by abundance of the marine dinocysts as Andalusiella polymorpha, Senegalinium bicavatum and Trichodinium castanea, Isabelidinium sp. Senegalinium granulostriatum, Senegalinium sp., Andalusiella sp., Deflandrea obliquipes, Senegalinium laevigatum, Batiacasphaera sp., Chytroeisphaeridia sp. with very poor land-derived miospores such as Verrucatosporites sp., Todisporites sp., Cyathidites australis, Laevigatosporites sp., Cyathidites minor, and Laricoidites magnus (Abu El Ella 2011). The abundance of marine dinocysts and poor land-derived miospores reflects deposition through offshore open marine conditions within the range of middle to outer neritic shelf environment.
Fig. 17

Examples of the open marine deposits of the Dakhla Formation, well BHK-5 (image data presented courtesy of SeaDragon-Egypt)

Paleogeographic evolution

To summarize the Cretaceous lithofacies and paleogeography, a variety of maps have been drawn based on boreholes data and outcrop sections. The paleogeographic maps are mainly based on the coarse/fine ratio contours and modified by the other parameters. The informal stages used for paleogeographic reconstruction considered in this paper are the Early Cretaceous Six Hills stage, Aptian Abu Ballas stage, Albian–Cenomanian Sabaya stage, Cenomanian–Turonian Taref-Maghrabi stage, Coniacian–Santonian Quseir stage, and Campanian–Maastrichtian Dakhla stage. The Six Hills stage is further subdivided into a six A–G stages, based on the 6 members of the Six Hills Formation. Chronological data in these formations are not sufficient to be certain that each of these maps (Fig. 18) accurately represents a time slice. We have followed Abu El Ella (2006, 2011) in addition to two main lithologic markers: the Coniacian–Santonian Coal bearing horizon and Top Barremian (top of G member of Six Hills Formation).
Fig. 18

Tectonic and sedimentary evolution of the Komombo Basin: ae Berriasian–Early Barremian. a Rift initiation systems tract. b Rift climax and ce late rifting progradation of braidplain deposits. f, g Late Barremian rift phase, f development of distal floodbasin or shallow lake and paleosols. g Late rifting stage with development of mid-/proximal braidplain. h, j, k Aptian–Albian rifting. lo Epicontinental sag and tectonic inversion

Basin evolution

The seven depositional systems described above represent the preserved Cretaceous sediments in the Komombo Basin. The basin evolution is characterized by successive tectonic rifting. The first phase begins in the Barrasian-Hautervian, the second phase begins during Late Barremian, and the third phase begins in the Aptian time. The first and third rifting phases are correlated with Early Cretaceous rifting phases which were recorded in North and Central Africa (Guiraud and Maurin 1992). The second Late Barremian rifting phase represents a transitional event between the two rifting phases.

The first phase of basin evolution is characterized by fluvial fan, lacustrine, and braidplain depositional systems in successive cycles or a general upward coarsening sequence (A–E members of the Six Hills Formation, Fig. 2b). The second phase, distal floodbasin and paleosols of F member of the Six Hills Formation are well developed and were followed by mid-/proximal braidplain of G member of the Six Hills Formation, generally in an upward coarsening sequence (Fig. 2b). The boundary between the first and second rifting phases is considered as unconformity due to reversing of the sediment dispersal trends (Figs. 8, 10). On the other hand, the boundary between the second and third rifting phases was characterized by marine ingression. A similar upward coarsening succession was described in the Muglad Basin (Central Sudan) but with a wider time span extending from Berriasian to Early Albian (McHargue et al. 1992). Numerous rift basins along the NE African margin were showing strong subsidence during Neocomian and Barremian times, notably the Alamein, Abu Gharadig, and Shushan in the northern Western Desert of Egypt (Bayoumi and Lotfy 1989; Taha 1992; Moustafa et al. 1998; Rodriguez 2015) and the Hameimat and Sarir in the southeastern Sirt (Exploration Staff of the Arabian Gulf Oil Company 1980).

During the third phase, lacustrine/lagoonal deposits (Abu Ballas Formation) are dominated by the tectonic subsidence and marine ingression. They are followed by the existence of fluvio-estuarine depositional system (Taref-Sabaya and Quseir formations). The tidal-flat and associated deposits were the main elements of the basin fill characterizing the third stage. The Aptian–Albian rifting phase is well documented in Abu Gabra, Anaza, Kaisur Basin of Sudan (Bosworth 1992), and Alamein and Abu Gharadig basins of North Western Desert (El Emam et al. 1990; Rodriguez 2015) in addition to Beni Suef Basin (Zahran et al. 2011). These rifting phases are recorded as well-defined distinct successions, here interpreted in terms of the external controls on the basin architecture (Fig. 18a–o). The successive rifting phases will be discussed in detail below.

Berriasian–Early Barremian rifting

Rift initiation

This rifting phase is almost synchronous with Late Berriasian continental rifting in NE Africa–Arabia (see Maurin and Guiraud 1990; Guiraud and Maurin 1991, 1992). The initial rifting stage is characterized by increasing rate of fault-related subsidence. The onset of deposition in the Komombo Basin is recorded in the A member of the Six Hills Formation being penetrated in the depocenter of the basin (Figs. 5, 6, 18) and is characterized by conglomerates, conglomeratic sandstones, and mudstones of the fluvial fan depositional system. This sedimentary package is characterized by wedge shape, which attains 75 m thickness and limited areal extent. Thus, I inclined to consider the deposition of this sedimentary package occurred in small fault-bounded basin (Fig. 18a). The small isolated depocenters, developed prior to the linkage of the early faults due to fault propagation (e.g., Gawthorpe and Leeder 2000), are diagnostic features of the rift initiation according to Prosser (1993). A similar succession was deposited at Sudan troughs such as Muglad Basin (McHargue et al. 1992).

This rift initiation is characterized by the establishment of locally sourced transverse and axial sedimentary lobes (Ravnås and Steel 1998). These are dominated by coarse-grained clastics of the stream-dominated fluvial fan in the axial domain and floodout-dominated fluvial fan in the lateral hanging wall. The hanging wall fluvial fan shows general progradation due to successive increase in the accommodation with the rotation of the hanging wall block (See Leeder and Gawthorpe 1987). This was shown from the general upward coarsening and thickening of the coarse-grained facies. The axial stream-dominated alluvial plain shows general aggradation to retrogradation due to the relative higher tectonic subsidence rate in contrary to the sediment supply rate (Fig. 4a).

Spatial distribution and provenance and paleocurrent patterns of the stream-dominated fluvial fan (Fig. 4a) point to an axial river system with transport directions from north to south and southwest with a northern source area (Fig. 18a). This phase of basin evolution is interpreted as the rift initiation with maximum preserved thickness of around 61 m.

Rift climax

The approximately 137-m-thick B member of the Six Hills Formation was probably deposited during this stage in a starved sedimentary environment where high rate of fault-related subsidence produced relatively deep anoxic water with relatively low clastic input. This resulted in deposition of offshore lacustrine mudstones (Fig. 9a, b). The limitation of the B member of the Six Hills Formation to the basin depocenter reflects the relatively high tectonic subsidence in the basinal area with more reduced sediment supply and limited drainage of the basin (Fig. 18b). Similar organic-rich lacustrine mudstones were deposited in Muglad Basin (McHargue et al. 1992) and Salamat trough of Central African Republic-Chad (Genik 1992). The transition from the fluvial fan system of A member of the Six Hills Formation to the lacustrine system of the B member delineates a transgressive surface.

Late rift

This late stage of Berriasian–Barremian rifting is characterized by gradual increase in the clastic supply relative to the tectonic subsidence with general successive progradation of the axial and lateral lobe/channel systems. The deposits of this stage show the braidplain progradation or basinward shift of the facies toward the floodbasin or shallow water lake. This will be encountered from the general upward coarsening pattern of the C–E members of the Six Hills Formation. This progradation is very clear from the vertical facies transition from distal to proximal braidplain (Fig. 18c–e). The proximal braidplain deposits are widely distributed through the rift basin especially nearby the master bounding fault of the Komombo Basin as seen from the drilled wells and the surface outcrops (Figs. 8a, 18e).

Spatial distribution and provenance and paleocurrent patterns of the mid- to proximal braidplain (Fig. 8a) refer to the complex channels system with transport directions from north/northwest to south and southeast with a northern source area (Fig. 18d, e).

Late Barremian rifting

The Late Barremian rifting phase represents the second phase of rifting at Komombo Basin and includes two stages as follow:

Rift climax

The synrift stage characterized by the dominance of the distal floodbasin/distal braidplain to shallow lacustrine mudstone (F member of the Six Hills Formation) reflects the increase in the rate of the accommodation outpaced the sediment supply (high-accommodation-system tract, Fig. 18f). The distal braidplain deposits characterize the underfilled phase in the evolution of the basin.

Late rift

This stage is characterized by tectonic sag as the distal floodbasin/distal alluvial plain succession passes upwards to mid- and proximal braidplain system with overbank deposits of G member of the Six Hills Formation (Figs. 10a, 18g) of the filling phase. The general change from underfilled to filled conditions is attributed to a shift in the balance between accommodation and the ability of sedimentary systems to fill the available space. The Late Barremian rifting phase shows reversing in the sediment transport direction from south/southeast to north/northwest with change in the source area to the south (Fig. 10a).

Aptian–Albian rifting

Rift climax

This synrift stage is characterized by marine flooding through the Aptian with dominance of the Abu Ballas mudstones that reflect the high rate of tectonic subsidence with low sediment supply rate. This event represents the first marine ingression in the Komombo Basin.

The dominance of the lacustrine/lagoonal deposits (Fig. 10a, b) reflects the creation of accommodation which outpaced the sediments supply from the hinterland source area. The lack of dinoflagellates and sparse pollens and spores reflects the deposition of lacustrine to lagoonal setting (Fig. 18h).

Late rifting

During the Sabaya stage, no new accommodation is generated and sediment supply gradually consumes the available space leading to a change from an underfilled to overfilled basin. Typically, sediment supply increases after a tectonic pulse, when the newly created differential relief between the source area and the basin allows for a more efficient delivery of sediment to the basin. The tectonic subsidence that was associated with deposition of Abu Ballas shales was followed by increase in the siliciclastics with dominance mixed-load fluvial channels during Sabaya-Taref Formation (Fig. 18j, k). The progradational pattern of siliciclastics through Abu Ballas to Sabaya-Taref stages probably reflects the waning of tectonic activity.

Post-rifting phase (epicontinental sag)

The rifting episode described above was followed during the Mid-Cretaceous by a tectonically quiet period (Guiraud 1998). This phase is characterized by the tectonic quiescence and degradation of source rock with dominance of fine-grained facies deposition with the development of the epicontinental sea. This stage began with major marine flooding of the Albian–Cenomanian systems with the dominance of the tidally affected fluvial channels, interfluve areas, and development of tidal estuaries (Figs. 11, 12, 13, 18k, l). This reflects the growth of the sea level role on the tectonic role in the deposition and sediment architectures and distribution. The gradual decrease in the clastic supply from the hinterland shows the dominance of the fine-grained deposition in mudflats and paleosol formation. The rate of relative sea level rise increased through Turonian–Santonian with development of the tidal flats and tidal channels in the inner shelf setting (Figs. 14, 18m). During the late Quseir (Santonian), the marine facies regression occurred with gradual decrease in the marine influences. This was followed by increase in the fluvial action and formation of tidally affected deltas (Figs. 14, 15, 16, 18n). This marine regression is confirmed by vertical facies transition from mudflats to prograding deltaic mouth bar deposits. This was eroded or capped by the fluvial or deltaic channels that intercalated with coaly and mudstone interbeds of the deltaplain setting. This event is probably simultaneous with Late Cretaceous Syrian Arc inversion pulses. Finally, the whole basin and surrounding areas were covered by Campanian–Maastrichtian (Dakhla) sea with deposition of open marine muds (Figs. 17a, b, 18o). The open marine depositional system was confirmed by the abundance of marine dinocysts and the limited land-derived miospores.

Conclusion

The Komombo Basin is an approximately 70-km-long and 30-km-wide, WNW-ESE-oriented Cretaceous rift basin formed during opening of the Neotethys. The sedimentary fill of the Komombo Basin can be ascribed to seven depositional systems: fluvial fan, braidplain, siliciclastic lakes, fluvio-estuarine, lake/lagoon, tidally affected deltas, and open marine settings. These depositional systems are in part coeval and in part succeed each other during the basin evolution.

In agreement with the main literature, the study reveals two main rifting phases commonly identified in the Central and North African Mesozoic basins through the Early Cretaceous. These two main phases of rifting are recognized in the history of Komombo Basin: Berriasian–Early Barremian and Aptian–Albian. A third additional Late Barremian rifting stage has been identified in the transition between those two main rifting phases. The Berriasian–Early Barremian rifting includes three stages: The rift initiation stage (A member of Six Hills Formation) is represented by conglomeratic and coarse-grained sandstones of fluvial fan-dominated successions. The rift climax stage (B member of the Six Hills Formation) of basin evolution began with a major reactivation tectonic subsidence with deposition in an anoxic lacustrine setting. This was followed by gradual decrease in tectonic subsidence and growth of the sediment supply to outpace the tectonic accommodation resulting in the development of the prograding distal to proximal braidplain deposits with late rift phase. The spatial distribution of the depositional systems and associated paleocurrent patterns indicate southerly sediment dispersal during Berriasian–Early Barremian. This rifting phase was described in many basins along the Northeastern African margin such as Hameimat and Sarir basins of Libya and Shushan and Abu Gharadig basins of Egypt that were dominated by fluvio-lacustrine deposits (Guiraud and Bosworth 1999) as well as Muglad Basin of Central Sudan (Mchargue et al. 1992).

Aptian–Albian rifting was associated with the creation of new accommodation space and accompanied marine ingression. This led to the formation of mixed-load fluvial channel and tidally affected fluvial channel with successive development of tidal estuaries (indicating the first ingression of the sea into the graben). This was followed by further transgression with the development of extensive tidal-flat domain with some tidal channels. This succession is interpreted as the record of the initial and continued thermal subsidence prior to post-rift thermal cooling, recorded in the Turonian–Santonian (Taref-Maghrabi stage). By the end of this stage, delta deposits are formed as consequence of marine regression and probably Santonian tectonic inversion. The post-rift deltaic deposits and the following open marine muds of Dakhla stage were deposited in a broader area than the synrift successions, recording deposition in a shallow sea-oriented parallel to the former rift. The Aptian–Albian rifting was described in Central African basins such as Abu Gabra, Anza, and Kaisur basins (Bosworth 1992). During the Albian–Aptian, most of Cretaceous rifts of southern Sudan represent sites for lacustrine sedimentation in the, resulting in widespread deposition of source rocks (Schull 1988). In the north Western Desert (Abu Gharadig Basin), the Dahab shales represent the deposition under tectonic subsidence (El Emam et al. 1990) as well as lower Kharita shales of Beni Suef Basin (Zahran et al. 2011).

An additional rifting stage has been identified in the transition between the two main phases: Berriasian–Early Barremian and Aptian–Albian. This stage (Late Barremian) is associated with creating accommodation with the abundance of distal floodbasin/distal braidplain deposits and paleosol formation. This was followed by a relative decrease in the tectonic subsidence and increase in the sediment supply from the hinterland with the development of mid- and proximal braidplain deposits. The spatial distribution of the depositional systems and associated paleocurrent patterns indicate reversing of sediment dispersal direction, formerly southerly during Berriasian–Early Barremian to Northerly during Late Barremian.

Notes

Acknowledgments

The author is grateful to SeaDragon-Egypt for the permission to use their data and to publish this paper. Prof. M. Darwish and Prof. A. El Manawi (Cairo University) and Eng. Alaa Ghoneimy (General Manager of south asset, SeaDragon-Egypt) are acknowledged for their critical and thorough reviews and suggestions. Dr. Ray Smith (CSIRO, Australia) is thanked for reviewing and editing the manuscript. The author is also grateful to the two reviewers: Prof. Thomas Voigt and Prof. Jochen Kuss, for their critical reading that improved the paper. The author also thanks Editor Prof. Wolf-Christian Dullo for handling the article for International Journal of Earth Sciences.

References

  1. Abu El Ella N (2006) Biostratigraphy and lithostratigraphy of Komombo-1, Komombo-2, Komombo-3, Nuqra-1 and Kharit-1 Wells, Upper Egypt, Earth Resources Exploration (EREX). Unpublished internal report, Cairo, 26 pGoogle Scholar
  2. Abu El Ella N (2011) Biostratigraphic studies for the AlBaraka oil field, Upper Egypt, Earth Resources Exploration (EREX). Unpublished internal report, Cairo, 250 pGoogle Scholar
  3. Almeida RP, Janikian L, Fragoso-Cesar ARS, Marconato A (2009) Evolution of a rift basin dominated by subaerial deposits: the Guaritas Rift, Early Cambrian, Southern Brazil. Sed Geol 217:30–51CrossRefGoogle Scholar
  4. Archangelsky S, Gamerro JC (1967a) Pollen grains found in coniferous cones from the Lower Cretaceous of Patagonia (Argentina). Rev Palaeobot Palynol 1:179–182CrossRefGoogle Scholar
  5. Archangelsky S, Gamerro JC (1967b) Spore and pollen types of the Lower Cretaceous in Patagonia (Argentina). Rev Palaeobot Palynol 1:211–218CrossRefGoogle Scholar
  6. Ashley GM (1990) Classification of large-scale subaqueous bedforms: a new look at an old problem. J Sediment Petrol 60:160–172CrossRefGoogle Scholar
  7. Awad MZ (1994) Stratigraphic, palynological and palaeoecological studies in the east-central Sudan (Khartoum and Kosti Basins), Late Jurassic to Mid-Tertiary. Berl Geowiss Abh 161:1–163Google Scholar
  8. Ayyad MH, Darwish M (1996) Syrian Arc structures: a unifying model of inverted basins and hydrocarbon occurrences in North Egypt. Egyptian General Petroleum Corporation Seminar, Cairo, 19 pGoogle Scholar
  9. Batten DJ, Uwins PJR (1985) Early-Late Cretaceous (Aptian–Cenomanian) Palynomorphs. J Br Micropalaeontol Soc 4(1):151–168CrossRefGoogle Scholar
  10. Bayoumi AI, Lotfy HI (1989) Modes of structural evolution of Abu Gharadig Basin, Western Desert of Egypt, as deduced from seismic data. J Afr Earth Sc 9:273–287CrossRefGoogle Scholar
  11. Birkeland PW (1984) Soils and geomorphology. Oxford University Press, New YorkGoogle Scholar
  12. Blair TC (1987) Tectonic and Hydrologic controls on cyclic alluvial fan, fluvial, and lacustrine rift-basin sedimentation, Jurassic-Lower Cretaceous Todos Santos Formation, Chiapas, Mexico. J Sediment Petrol 57:845–862Google Scholar
  13. Boltenhagen E (1977) Microplancton du crétacé supérieur du Gabon. Cah Paleontol 4:1–150Google Scholar
  14. Bosworth W (1992) Mesozoic and early Tertiary rift tectonics in East Africa. Tectonophysics 209:115–137CrossRefGoogle Scholar
  15. Bosworth W, El-Hawat AS, Helgeson DA, Burke K (2008) Cyrenaican “shock absorber” and associated inversion strain shadow in the collision zone of northeast Africa. Geology 36:695–698CrossRefGoogle Scholar
  16. Brenner GJ (1968) Middle Cretaceous spores and from northeastern Peru. Pollen Spores 10:341–382Google Scholar
  17. Bridges PH, Leeder MR (1976) Sedimentary model for intertidal mudflat channels, with examples from the Solway Firth, Scotland. Sedimentology 23:533–552CrossRefGoogle Scholar
  18. Bumby AJ, Guiraud R (2005) The geodynamic setting of the Phanerozoic basins of Africa. J Afr Earth Sci 43:1–12CrossRefGoogle Scholar
  19. Burke KC, Dewey JF (1974) Two plates in Africa during the Cretaceous? Nature 249:313–316CrossRefGoogle Scholar
  20. Cain SA, Mountney NP (2009) Spatial and temporal evolution of a terminal fluvial fan system: the Permian Organ Rock Formation, South-east Utah, USA. Sedimentology 56:1774–1800CrossRefGoogle Scholar
  21. Choi KS, Dalrymple RW, Chun SS, Kim SP (2004) Sedimentology of modern inclined heterolithic stratification (IHS) in the macrotidal Han River delta, Korea. J Sediment Res 74:677–689CrossRefGoogle Scholar
  22. Crossley R (1984) Controls of sedimentation in the Malawi rift valley, central Africa. Sed Geol 40:33–50CrossRefGoogle Scholar
  23. Dalrymple RW (1992) Tidal depositional systems. In: Walker RG, James NP (eds) Facies models and sea level changes. Geological Association of Canada, St. John’s, pp 195–218Google Scholar
  24. Dalrymple RW, Choi K (2007) Morphologic and facies trends through the fluvial–marine transition in tide-dominated depositional systems: a schematic framework for environmental and sequence-stratigraphic interpretation. Earth Sci Rev 81:135–174CrossRefGoogle Scholar
  25. Dalrymple RW, Baker EK, Harris PT, Hughes M (2003) Sedimentology and stratigraphy of a tide-dominated, foreland-basin delta (Fly River, Papua New Guinea). In: Sidi FH, Nummedal D, Imbert P, Darman H, Posamentier HW (eds) Tropical deltas of Southeast Asia—sedimentology, stratigraphy and petroleum geology. SEPM Special Publication 76, pp 147–173Google Scholar
  26. De Mowbray T (1983) The genesis of lateral accretion deposits in recent intertidal rnudflat channels, Solway Firth, Scotland. Sedimentology 30:425–435CrossRefGoogle Scholar
  27. Dolson JC, Shann MV, Matbouly S, Harwood C, Rashed R, Hammouda H (2001) The petroleum potential of Egypt. In: MW Downey, JC Threet, WA Morgan (eds) Petroleum provinces of the 21st century, Tulsa, Oklahoma. AAPG Memoir, pp 453–482Google Scholar
  28. Dolson JC, Atta M, Blanchard D, Sehim A, Villinski J, Loutit T, Romine K (2014) Egypt’s future petroleum resources: a revised look into the 21st century. In: Marlow L, Kendall C, Yose L (eds) Petroleum systems of the Tethyan region. AAPG Memoir 106, pp 143–178Google Scholar
  29. Duchaufour P (1982) Pedology. Allen and Unwin, LondonCrossRefGoogle Scholar
  30. Eberth DA (1996) Origin and significance of mud-filled incised valleys (Upper Cretaceous) in southern Alberta, Canada. Sedimentology 43:459–477CrossRefGoogle Scholar
  31. El Beialy SY, Al Hitmi H (1994) Micropaleontology and palynology of the Lower and Middle Cretaceous Thamama and Wasia groups, DK-C well, Dukhan Oil Field, Western Qatar. Sci Géol Bull 47(1–4):67–95Google Scholar
  32. El Emam A, Bishop D, Dunderale I (1990) The hydrocarbon potential of the west Gindi area, Western Desert, Egypt. In: 10th Egyptian General Petroleum Cooperation, exploration and production seminar, Cairo, 14 pGoogle Scholar
  33. Exploration Staff of the Arabian Gulf Oil Company (1980) Geology of a stratigraphic giant-the Messlah oil field. In: Salem MJ, Busrewil MT (eds) The geology of Libya, vol 2. Academic Press, London, pp 521–538Google Scholar
  34. Fathy M, Salvadori L, Roberts G, Gouda MA (2010) Komombo: a new oil province in Upper Egypt (abst.). AAPG GEO 2012 Middle EastGoogle Scholar
  35. Fielding CR (1987) Coal depositional models for deltaic and alluvial plain sequences. Geology 15:661–664CrossRefGoogle Scholar
  36. Fischbein SA, Joeckel RM, Fielding CR (2009) Fluvial-estuarine reinterpretation of large, isolated sandstone bodies in epicontinental cyclothems, Upper Pennsylvanian, northern Midcontinent, USA, and their significance for understanding late Paleozoic sea-level fluctuations. Sed Geol 216:15–28CrossRefGoogle Scholar
  37. Frostick LE, Reid I (1987) Tectonic controls of desert sediments in rift basins ancient and modern. In: Frostick LE, Reid I (eds) Desert sediments: ancient and modern. Geological Society Special Publication 35, pp 53–68Google Scholar
  38. Gawthorpe RL, Leeder MR (2000) Tectono-sedimentary evolution of active extensional basins. Basin Res 12:195–218CrossRefGoogle Scholar
  39. Gawthorpe RL, Fraser AJ, Collier REL (1994) Sequence stratigraphy in active extensional basins: implications for the interpretation of ancient basin-fills. Mar Pet Geol 11:642–658CrossRefGoogle Scholar
  40. Genik GJ (1992) Regional framework, structural and petroleum aspects of rift basins in Niger, Chad and the Central African Republic (C.A.R.). Tectonophysics 213:169–185CrossRefGoogle Scholar
  41. Gübeli AA, Hochuli PA, Wildi W (1984) Lower Cretaceous turbiditic sediments from the Rif chain (northern Morocco)-palynology, strattgraphy and palaeogeographic setting. Geol Rundsch 73:1081–1114CrossRefGoogle Scholar
  42. Guiraud R (1998) Mesozoic rifting and basin inversion along the northern African Tethyan margin: an overview. In: Macgregor DS et al (eds) Petroleum geology of North Africa. Geological Society of London Special Publication 132, pp 217–230Google Scholar
  43. Guiraud R, Bellion Y (1995) Late Carboniferous to recent geodynamic evolution of the West Gondwanian cratonic Tethyan margins. In: Nairn A, Ricou LE, Vrielynck B, Dercourt J (eds) The ocean basins and margins 8, the Tethys Ocean. Plenum Press, New York, pp 101–124Google Scholar
  44. Guiraud R, Bosworth W (1997) Senonian basin inversion and rejuvenation of rifting in Africa and Arabia: synthesis and implications to platescale tectonics. Tectonophysics 282:39–82CrossRefGoogle Scholar
  45. Guiraud R, Bosworth W (1999) Phanerozoic geodynamic evolution of northeastern Africa and the northwestern Arabian platform. Tectonophysics 315:73–108CrossRefGoogle Scholar
  46. Guiraud R, Maurin JC (1991) Le rifting en Afrique au Crétacé inférieur: synthèse structurale, mise en évidence de deux étapes dans la genése des bassins, relations avec les ouvertures océaniques périafricaines. Bull Soc Géol Fr 162:811–823CrossRefGoogle Scholar
  47. Guiraud R, Maurin JC (1992) Early Cretaceous rifts of Western and Central Africa: an overview. Tectonophysics 213:153–168CrossRefGoogle Scholar
  48. Guiraud R, Bosworth W, Thierry J, Delplanque A (2005) Phanerozoic geological evolution of Northern and Central Africa: an overview. In: Catuneanu O et al (eds) Phanerozoic evolution of Africa. J Afr Earth Sci 43:83–143Google Scholar
  49. Herngreen GFW (1975) Palynology of middle and upper Cretaceous strata in Brazil. Mededelingen/Rijks Geologische Dienst, NS 26(3):39–91Google Scholar
  50. Ibrahim MIA (1996) Aptian-Turonian palynology of the Ghazalat-1 well (GTX-1), Qattara Depression, Egypt. Rev Paleobot Palynol 94:137–168CrossRefGoogle Scholar
  51. Ibrahim MIA, Abdel-Kireem MR (1997) Late Cretaceous palynofloras and foraminifera from Ain El-Wadi area, Farafra Oasis, Egypt. Cretac Res 18:633–660CrossRefGoogle Scholar
  52. Ito M, Matsukawa M, Saito T, Nichols ND (2006) Facies architecture and paleohydrology of a synrift succession in the Early Cretaceous Choyr Basin, southeastern Mongolia. Sed Geol 27:226–240Google Scholar
  53. Jain KP, Millepied P (1973) Cretaceous microplankton from Senegal Basin, N. W. Africa. 1. Some new genera species and combinations of dinoflagellates. The. Palaeobotanist 20(1):22–32Google Scholar
  54. Jenkins DA (1990) North and central Sinai. In: Said R (ed) Geology of Egypt. A.A. Balkema, Rotterdam, pp 361–380Google Scholar
  55. Jo HR, Rhee CW, Chough SK (1997) Distinctive characteristics of a streamflow-dominated alluvial fan deposit: Sanghori area, Kyongsang Basin (Early Cretaceous), southeastern Korea. Sed Geol 110:51–79CrossRefGoogle Scholar
  56. Kamel H (1990) Gravity map. In: Said R (ed) Geology of Egypt. A.A. Balkema Publishers, Rotterdam, pp 45–50Google Scholar
  57. Kampf N, Schwertmann U (1982) Goethite and hematite in a climosequence in southern Brazil and their application in classification of kaolinitic soils. Geoderma 29:27–39CrossRefGoogle Scholar
  58. Kaska HV (1989) A spore and pollen zonation of Early Cretaceous to Tertiary nonmarine sediments of central Sudan. Palynology 13:79–90CrossRefGoogle Scholar
  59. Klein GD (1971) A sedimentary model for determining paleotidal range. Geol Soc Am Bull 82:2585–2592CrossRefGoogle Scholar
  60. Klitzsch EH, Squyres CH (1990) Paleozoic and Mesozoic Geological history of Northeastern Africa based on new interpretation of Nubian strata. AAPG Bull 74(8):1203–1211Google Scholar
  61. Kusznir NJ, Marsden G, Egan SS (1991) A flexural cantilever simple shear/pure shear model of continental lithosphere extension: applications to the Jeanne d’Arc basin, Grand Banks, and Viking Graben, North Sea. In: Roberts AM, Yielding G, Freeman B (eds) The geometry of normal faults. Geological Society Special Publication 56, pp 41–60Google Scholar
  62. Lambiase JJ (1990) A model for tectonic control of lacustrine stratigraphic sequences in continental rift basins. In: BJ Katz (ed) Lacustrine basin exploration: case studies and modern analogs: AAPG Memoir 50, pp 265–276Google Scholar
  63. Lawal O, Moullade M(1986) Palynological biostratigraphy of Cretaceous sediments in the Upper Benue Basin, N.E. Nigeria (1). Revue de Micropaléontologie 29:61–83Google Scholar
  64. Leeder MR (1995) Continental rifts and proto-oceanic rift troughs. In: Busby CJ, Ingersoll RV (eds) Tectonics of sedimentary basins. Blackwell Science, Oxford, pp 119–148Google Scholar
  65. Leeder MR, Gawthorpe RL (1987) Sedimentary models for extensional tilt-block/half-graben basin. In: Coward MP, Dewey JF, Hancock PL (eds) Continental extensional tectonics. Geological Society Special Publication 28, pp 139–152Google Scholar
  66. Lister GS, Etheridge MA, Symonds PA (1986) Detachment faulting and the evolution of passive continental margins. Geology 14:246–250CrossRefGoogle Scholar
  67. Mahmoud MS, Deaf AS (2007) Cretaceous palynology (spores, pollen and dinoflagellate cysts) of the Siqeifa 1-x borehole, northern Egypt. Riv Ital Paleontol Stratigr 113:203–221Google Scholar
  68. Maizels J (1993) Lithofacies variations within sandur deposits: the role of runoff regime, flow dynamics and sediment supply characteristics. Sed Geol 85:299–325CrossRefGoogle Scholar
  69. Malloy RE (1972) An Upper Cretaceous dinoflagellate cyst lineage from Gabon, West Africa. Geosci Man 4:57–65CrossRefGoogle Scholar
  70. Mángano MG, Buatois LA, Aceñolaza GF (1996a) Trace fossils and sedimentary facies from an Early Ordovician tide-dominated shelf (Santa Rosita Formation, northwest Argentina)—implications for ichnofacies models of shallow marine successions. Ichnos 5:53–88CrossRefGoogle Scholar
  71. Mángano MG, Buatois LA, Maples CG, West R (1996b) Trace fossils from an Upper Carboniferous tidal shoreline (Stull Shale Member of eastern Kansas). In: 30th international geological congress, Beijing, Abstract 2, p 133Google Scholar
  72. Maurin JC, Guiraud R (1990) Relationships between tectonics and sedimentation in the Barremo-Aptian intracontinental basins of Northern Cameroon. In: Kogbe CA, Lang J (eds) African continental phanerozoic sediments. J Afr Earth Sci 10:331–340Google Scholar
  73. McHargue TR, Heidrick TL, Livingston JE (1992) Tectonostratigraphic development of the interior Sudan rifts, Central Africa. In: Ziegler PA (eds) Geodynamics of rifting, vol II. Case history studies on rifts: North and South America and Africa. Tectonophysics 213:187–202Google Scholar
  74. McKenzie D (1978) Some remarks on the development of sedimentary basins. Earth Planet Sci Lett 40:25–32CrossRefGoogle Scholar
  75. Mohsen SA (1992) Cretaceous plant microfossils from the subsurface of Kharga Oasis, Western Desert, Egypt. J Afr Earth Sci 14(4):567–577CrossRefGoogle Scholar
  76. Morley CK (1989) Extension, detachments, and sedimentation in continental rifts (with particular reference to East Africa). Tectonics 8:1175–1192CrossRefGoogle Scholar
  77. Mossop GD, Flach PD (1983) Deep channel sedimentation in the Lower Cretaceous McMurray Formation, Athabasca Oil sands, Alberta. Sedimentology 30:493–509CrossRefGoogle Scholar
  78. Moustafa AR, Khalil MH (1990) Structural characteristics and tectonic evolution of north Sinai fold belts. In: Said R (ed) Geology of Egypt. A.A. Balkema, Rotterdam, pp 381–389Google Scholar
  79. Moustafa AR, El-Badrawy R, Gibali H (1998) Pervasive E-ENE oriented faults in northern Egypt and their effect on the development and inversion of prolific sedimentary basins. In: Proceedings of the of 14th petroleum conference, vol 1. Egyptian General Petroleum Corporation, Cairo, pp 51–67Google Scholar
  80. Nagati M (1986) Possible Mesozoic rifts in Upper Egypt: an analogy with the geology of Yemen-Somalia rift basins. In: Proceeding of the 8th petroleum exploration conference, vol 2. Egyptian General Petroleum Corporation, Cairo, pp 205–231Google Scholar
  81. Nio SD, Yang CS (1991) Diagnostic attributes of clastic tidal deposits: a review. In: Smith DG, Reinson GE, Zaitlin BA, Rahmani RA (eds) Clastic tidal sedimentology. Canadian Society of Petroleum Geologists Memoir 16, pp 3–28Google Scholar
  82. Prosser S (1993) Rift-related linked depositional systems and their seismic expression. In: Wouldiams GD, Dobb A (eds) Tectonics and seismic sequence stratigraphy. Geological Society Special Publication 71, pp 35–66Google Scholar
  83. Rauscher R, Doubinger J (1982) Les dinokystesdu Maestrichtien Phosphate Maroc. Sci Géol Bull 35(3):97–116Google Scholar
  84. Ravnås R, Steel RJ (1998) Architecture of Marine Rift-Basin Successions. AAPG Bull 82:110–146Google Scholar
  85. Regali MSP (1989) Evoluçao da paleoflora no Cretaceo das margens equatorial de nordeste do Brasil. Rev Esc Minas 42:17–33Google Scholar
  86. Reineck HE (1958) Longitudinale schrägschicht in Watt. Geol Rundsch 47:73–82CrossRefGoogle Scholar
  87. Reineck HE, Wunderlich F (1968) Classification and origin of flaser and lenticular bedding. Sedimentology 11:99–104CrossRefGoogle Scholar
  88. Reyre Y (1973) Palynologie du Mésozoique Saharien. Mèmoires, Museum National Historie Naturelle, N. S., 27:1–284Google Scholar
  89. Rodriguez P (2015) The Mesozoic rifting of Alamein basin (Western Desert, Egypt): 3D evidence of transtension in the southern Tethyan margin. European Regional Conference and Exhibition, Lisbon, Portugal. Abstract. AAPG Search and Discovery Article #90226. http://www.searchanddiscovery.com/abstracts/html
  90. Rosendahl BR (1987) Architecture of continental rifts with special reference to east Africa. Annu Rev Earth Planet Sci Lett 15:445–503CrossRefGoogle Scholar
  91. Said R (1990) Mesozoic. In: Said R (ed) Geology of Egypt. Balkema, Rotterdam, pp 451–486Google Scholar
  92. Schrank E (1987) Paleozoic and Mesozoic palynomorphs from northeast Africa (Egypt and Sudan) with special reference to Late Cretaceous pollen and dinoflagellates. Berl Geowiss Abh Reihe A 75(1):249–310Google Scholar
  93. Schrank E (1991) Mesozoic palynology and continental sediments in NE Africa (Egypt and Sudan)—a review. J Afr Earth Sci 12:363–373CrossRefGoogle Scholar
  94. Schrank E (1992) Nonmarine Cretaceous correlations in Egypt and northern Sudan: palynological and palaeobotanical evidence. Cretac Res 13:351–368CrossRefGoogle Scholar
  95. Schrank E, Ibrahim MIA (1995) Cretaceous (Aptian–Maastrichtian) palynology of foraminifera-dated wells (KRM-1, AG-18) in northwestern, Egypt, vol 177. Berliner Geowissenschaftliche Abhandlungen, pp 1–44Google Scholar
  96. Schrank E, Mahmoud MS (1998) Palynology (pollen, spores and dinoflagellates) and Cretaceous stratigraphy of the Dakhla Oasis, Central Egypt. J Afr Earth Sc 26:167–193CrossRefGoogle Scholar
  97. Schull TJ (1988) Rift basins of interior Sudan, Petroleum exploration and discovery. AAPG Bull 27:1128–1142Google Scholar
  98. Sehim A (1993) Cretaceous tectonics in Egypt. Egypt J Geol 37:335–372Google Scholar
  99. Smith DG (1987) Meandering river point bar lithofacies models: modern and ancient examples compared. In: Ethridge FG, Flores RM, Harvey MD (eds) Recent developments in fluvial sedimentology. SEPM Special Publication 39, pp 83–91Google Scholar
  100. Smith DG (1988) Modern point bar deposits analogous to the Athabasca Oil Sands, Alberta, Canada. In: de Boer PL, van Gelder A, Nio SD (eds) Tide-influenced sedimentary environments and facies. Reidel Publishing Company, Dordrecht, pp 417–432CrossRefGoogle Scholar
  101. Smith DG, Reinson GE, Zaitlin BA, Rahmani RA (1991) Clastic tidal sedimentology. Canadian Society of Petroleum Geologists Memoir 16, 387 ppGoogle Scholar
  102. Steel RJ (1988) Coarsening-upward and skewed fan bodies: symptoms of strike-slip and transfer fault movement in sedimentary basins. In: Nemec W, Steel RJ (eds) Fan deltas: sedimentology and tectonic settings. Blackie, Glasgow, pp 75–83Google Scholar
  103. Surlyk F (1978) Submarine fan sedimentation along fault scarps on tilted fault blocks (Jurassic–Cretaceous boundary, East Greenland): Grønlands Geologiske Undersøgelse. Bulletin 128:1–108Google Scholar
  104. Surlyk F (1989) Mid-Mesozoic synrift turbidite systems: controls and predictions. In: Collinson JD (ed) Correlation in hydrocarbon exploration. Norwegian Petroleum Society, pp 231–241Google Scholar
  105. Taha MA (1992) Mesozoic rift basins in Egypt: their southern extension and impact on future exploration. In: Abdine S (ed) Proceedings of the 11th petroleum exploration and production conference. Egyptian General Petroleum Corporation, Cairo, Egypt, pp 1–19Google Scholar
  106. Thomas RD, Smith DG, Wood JM, Visser J, Calverly-Range EA, Koster EH (1987) Inclined heteolithic stratification terminology; description, interpretation and significance. Sed Geol 53:123–179CrossRefGoogle Scholar
  107. Todd SP (1989) Stream-driven, high-density gravelly traction carpets: possible deposits in the Trabeg conglomerate Formation, SW Ireland and theoretical considerations of their origin. Sedimentology 36:513–530CrossRefGoogle Scholar
  108. Tooth S (2000a) Downstream changes in dryland river channels: the Northern Plains of arid central Australia. Geomorphology 34:33–54CrossRefGoogle Scholar
  109. Tooth S (2000b) Process, form and change in dryland rivers: a review of recent research. Earth Sci Rev 51:67–107CrossRefGoogle Scholar
  110. Tooth S (2005) Splay formation along the lower reaches of ephemeral rivers on the Northern Plains of central Australia. J Sediment Res 75:636–649CrossRefGoogle Scholar
  111. Van den Berg JH, Boersma JR, van Gelder A (2007) Diagnostic sedimentary structures of the fluvial-tidal transition zone—evidence from deposits of the Rhine and Meuse. Neth J Geosci 86:287–306Google Scholar
  112. Visser MJ (1980) Neap-spring cycles reflected in Holocene subtidal large-scale bedform deposits: a preliminary note. Geology 8:543–546CrossRefGoogle Scholar
  113. Wernicke B (1985) Uniform sense normal simple shear of the continental lithosphere. Can J Earth Sci 22:108–125CrossRefGoogle Scholar
  114. Wood B, Zakariya A, Hady AA (2012) Resetting the geological framework of the Al Baraka field, Komombo Concession, Upper Egypt. Abstract. In: 10th Middle East Geosciences Conference and Exhibition. Manama, Bahrain. AAPG Search and Discovery Article #90141. http://www.searchanddiscovery.com/abstracts/html
  115. Wycisk P (1994) Correlation of the major late Jurassic–early Tertiary low- and highstand cycles of south-west Egypt and north-west Sudan. Geol Rundsch 83:759–772CrossRefGoogle Scholar
  116. Zahran H, Abu Elyazid K, Mohamad M (2011) Beni Suef Basin: the key for exploration future success in Upper Egypt. AAPG annual convention and exhibition, Houston, USA. Search and discovery article #10351. http://www.searchanddiscovery.com/documents
  117. Ziegler PA (1992) Geodynamics of rifting and implications for hydrocarbon habitat. Tectonophysics 215:221–253CrossRefGoogle Scholar

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Authors and Affiliations

  1. 1.Department of Geology, Faculty of ScienceCairo UniversityGizaEgypt

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