Facies and architecture of a coarse-grained alluvial-dominated incised valley fill: a case study from the Oligocene Gebel Ahmar Formation, Southern Tethyan-shelf (northern Egypt)

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Proceedings of the Geologists’ Association 128 (2017) 234–255

Contents lists available at ScienceDirect

Proceedings of the Geologists’ Association journal homepage: www.elsevier.com/locate/pgeola

Facies and architecture of a coarse-grained alluvial-dominated incised valley ﬁll: a case study from the Oligocene Gebel Ahmar Formation, Southern Tethyan-shelf (northern Egypt) S.S. Selim Department of Geology, Faculty of Science, Cairo University, Giza, 12613, Egypt

A R T I C L E I N F O

Article history: Received 26 June 2016 Received in revised form 1 December 2016 Accepted 15 December 2016 Available online 9 February 2017 Keywords: Incised valley Braid plain Oligocene Gulf of Suez Gebel Ahmar Formation

A B S T R A C T

Exposures of multistorey, alluvial deposits from the Oligocene Gebel Ahmar Formation in the Cairo-Suez province (north Eastern Desert, Egypt) show the architecture of an up to 35 m thick continuously prograding ﬂuvial/alluvial ﬁlling of an incised valley. The Oligocene base level fall resulted in cannibalization of the Eocene bedrock. Subsequent baselevel rise created accommodation space that was ﬁlled by deposition of four stacked storeys: lower storeys (1-2) of low sinuosity sandy braid plains and upper storeys (3-4) of gravelly braid plain. These braid plains were sourced from exposed Upper Cretaceous-Eocene and Paleozoic-Lower Cretaceous siliciclastic successions to the south. These successions dominate the Galala-Araba inverted structures. The sandy braid plain channel belts mainly downstream accretion (DA), downstream oblique accretion(DLA), lateral accretion (LA), sandy bedforms (SB), channel (CH), and Hollow (HO) elements, while the gravelly braid plain consists mainly of gravel bars and sheets (GB), gravel-sandstone foresets (GSF), gravel-sand couplets (GSC), and scour pool ﬁlling (SPF) architectures. Incised valley incision is potentially linked to a global drop of sea level caused by glaciation, although hinterland tectonism (i.e. Late Cretaceous-Paleogene tectonic inversion and Late Eocene-Oligocene crustal updoming in the source terrains) as well as Late Oligocene-Miocene rifting play a signiﬁcant role in the subsequent ﬁlling. The hinterland tectonism as well as the climate controls the sediment supply. The understanding of the nature of the Oligocene incised valley ﬁll helps in the constrain potential down depositional dip hydrocarbon reservoirs in Nile Delta, East Mediterranean basins, and similar settings in passive continental margins. © 2016 The Geologists' Association. Published by Elsevier Ltd. All rights reserved.

1. Introduction The incised valley systems are of great interest due to their ﬁlling reﬂects changes in response to climate, sea-level, tectonic, and anthropic inﬂuences. Incised valleys are being critical in sequence stratigraphy, and wetland management (Dalrymple et al., 1994; Rodriguez et al., 1998; Blum and Aslan, 2006; Mattheus and Rodriguez, 2011; Srivastava et al., 2013; Breda et al., 2016; Main et al., 2016). In addition, incised valleys can be key for hydrocarbon reservoirs (Martin et al., 2011; Mattheus and Rodriguez, 2011). The majority of incised valleys preserved in the stratigraphic record are thought to be cut and ﬁlled in response to a fall and subsequent rise of relative sea level (Boyd et al., 2006). Incision and ﬁlling of an incised valley can be driven also by factors not related to relative sea-level changes, such as variations in ﬂuvial discharge due to

E-mail address: [email protected] (S.S. Selim).

climatic and/or tectonic changes (Schumm et al., 1987; Blum, 1992; Schumm, 1993; Holbrook, 2001; Aldinucci et al., 2007; Rasmussen, 2014). These incised valleys are dominated by ﬂuvial sandstone ﬁll with ﬁne-grained parting, commonly cited in the literature (Wright and Marriot, 1993; Aitken and Flint, 1994; Willis, 1997; Blum and Price, 1998; Arnott et al., 2000; Aldinucci et al., 2007; Wang et al., 2015; Rasmussen, 2014; Bianchi et al., 2015, and references therein) but the documentation of the gravelly type that lacking the ﬁne-grained parting are less common. This case study offers an opportunity to study architecture and evolution of a coarse-grained alluvial incised valley ﬁll from the rock record and possible controlling factors such as tectonics and climate. This study focuses on the coarse-grained alluvial-dominated, valley ﬁll deposits of the Oligocene Gebel Ahmar Formation that overlies different levels of Eocene carbonate bedrock and underlies Miocene basalt sheets at Cairo-Suez province, northern Egypt (Fig. 1A). Surprisingly, no architectural studies have been carried out for this succession in NE Egypt. The incised valley setting is

http://dx.doi.org/10.1016/j.pgeola.2016.12.006 0016-7878/© 2016 The Geologists' Association. Published by Elsevier Ltd. All rights reserved.

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inferred due to its wedge-shape external cross-sectional geometry, basal unconformity surface, multistorey internal architectures, lithofacies, and stratigraphic relationships of this formation. By the Earliest Oligocene, deep subaerial valleys were incised into the exposed carbonate shelf (Dolson et al., 2005). These valleys incised during general erosional phase that characterized the Oligocene eustatic sea level drop, due to ice volume and cooling event (Miller et al., 2008). Mess et al. (2001) revealed a global climatic change toward cool conditions and development of large Antarctic ice sheets with shifting of vegetation from dominantly tropical to subtropical Eocene forests to more “mixed mesophytic forests”. The aim of this study is to describe and interpret the sedimentary facies, their origin and evolution, and to examine their architectural elements of barform and channel-inﬁll dimensions through detailed ﬁeld-based examinations of Oligocene Gebel Ahmar Formation. Using the sedimentological interpretations to document spatial and temporal variability in alluvial architecture, and discuss and integrate the implications of these ﬁndings as they relate to possible controls on valley incision and ﬁlling. The depositional architecture and stacking pattern of such deposits reﬂect ﬂuctuations in intrinsic parameters such as discharge and sediment supply, and accommodation (interaction between climate and tectonism). It is hoped that by gaining a good understanding of the incised valley ﬁll an insight into the down dip hydrocarbon reservoirs in the nearby east Nile Delta can be gained. The Oligocene reservoirs of East Nile Delta represent one of the most promising and future exploration targets (see El Naggar and El Morshedy, 2013; Selim, 2016). 2. Geological background

Fig. 1. (A) Paleogeography of the Oligocene strata along the northern rim of the Arabian-African plates, summarized after Rögl (1998), Meulenkamp and Sissingh (2003), Dolson et al. (2005),Issawi and McCauley (1993), Selim (2016). The major Oligocene to Miocene ﬂuvial incisions are highlighted by canyon-drainages. The inset map is Oligocene paleogeographic map (after Meulenkamp and Sissingh, 2003 ). (B) Geological map of the north Eastern Desert (after Klitzsch et al., 1987). The

Strong tectonic activity and rapid and intense changes in global climate occurred during the Oligocene (see Guiraud et al., 2005). The dramatic drop in the global sea level that occurred in the earliest Oligocene caused by a glacial episode (Miller et al., 2008), resulted in the emergence of most of the continental shelves around the world (Guiraud et al., 2005). By the end of Eocene, the NE Africa continental margin was characterized by a carbonate shelf that has been incised and extensively eroded with many northward-trending incised valleys (Dolson et al., 2014). These valleys were ﬁlled mainly with thick non-marine succession across northern Egypt (Dolson et al., 2014). The Oligocene Gebel Ahmar Formation at the Cairo-Suez province represents one of these valley ﬁlls. The Cairo-Suez province (Study area) is delimited to the south by Galala-Araba uplifted blocks and to the north by the Nile Delta (Fig. 1B). Galala-Araba uplifted blocks represent a part from the Syrian Arc (Kuss et al., 2000; Hussein and Abd-Allah, 2001), which display large NE-SW oriented folds in northern Egypt and Eastern Mediterranean regions, due to a convergence between Africa and Eurasia and the closure of the Neotethys during the Late Cretaceous-Paleogene (Moustafa and Khalil, 1995; Abd-Allah, 2008). Further south at NE corner of Afro-Arabian plate, mantle plume of Afar uplifted during the Late Eocene-Oligocene (Hofmann et al., 1997). To the west and southwest, NW-oriented faults were formed with the opening of the Gulf of Suez during Late OligoceneMiocene. The Gulf of Suez rift represents a continuation of the Red Sea rifting with dominant NW-oriented extensional faults, following the divergence of African plate away from Arabian plate (Patton et al., 1994 and Bosworth and McClay, 2001). Oligocene-aged outcrops in Egypt (Fig. 1) are sparse as a result of wide-scale erosion due to tectonically induced uplift, enhanced by a major and global mid-Oligocene drop in sea-level (Haq et al., 1988; Haq and Al-Qahtani, 2005). In the North Western Desert structure of the Cairo-Suez province after Hussein and Abd-Allah (2001). (C) Satellite image of the study area showing location of the studied proﬁles.

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Fig. 2. (A) Cross section through the Kattamiya-Sukhna road showing the geometry of the studied deposits ascribed to the valley ﬁll (not to scale). Note the basal boundary of the incised valley (dashed line) is not always observed, (B) The Miocene basalt sheet overlies the Oligocene Gebel Ahmar Fomation, (C) The sequence boundary between Upper Eocene bedrock and Oligocene valley ﬁll.

(Fig. 3), the lower Oligocene Gebel Qatrani Formation is composed of up to 340 m of estuarine and ﬂuvio-estuarine sandstones and sandy mudstones (Beadnell, 1905). In the Cairo-Suez province, the Oligocene Gebel Ahmar Formation is composed of ﬂuvial sandstones and conglomerates (Sadek, 1926; Shukri, 1954; Said, 1981). Down depositional dip in the Nile Delta, the Oligocene Dabaa Formation is composed of shelf and pelagic shales with sandstone interbeds (Dolson et al., 2002).

The Oligocene Gebel Ahmar Formation overlies different levels of Eocene carbonates and underlies Early Miocene ﬂood basalt sheets (Figs. 2 and 3 ), that followed by Early Miocene marine mixed siliciclastic and carbonate of Hommath Formation (Ismail and Abdel Ghany, 1999). The basal contact between the Oligocene and underlying Eocene bedrock is unconformable and is covered by basal lags (Fig. 2C). The Eocene bedrock comprises three formations from base to top; the basal shelfal limestone and

Fig. 3. The stratigraphic framework of the Oligocene of northern Egypt with data from different authors. The stratigraphic of Gebel Ahmar Formation in the study area is represented by the stratigraphic proﬁles.

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Fig. 4. Lithofacies photographs: (A) massive conglomerate (Gm), B) Stratiﬁed conglomerate (Gh) are capped by cross-stratiﬁed sandstone (Sp) with vertical and inclined skolithos ichnofossil, C) Cross-stratiﬁed sand-gravel clinoforms, D) Trough cross-stratiﬁed sand-gravel, E) trough cross-stratiﬁed sandstone, F) Convoluted cross-stratiﬁed sandstone (Sd).

dolomitic limestone of Qurn Formation (L. Bartonian) to lagoonal/ tidal carbonates of Wadi Garawi Formation (U. Bartonian), lagoonal marls and sandy limestone of Wadi Hof Formation (Upper Eocene, Strougo and Abd-Allah, 1990; Strougo et al., 1992). The Oligocene Gebel Ahmar Formation wedged out and is completely missed eastward in the Cairo-Suez province, as the Lower MioceneHommath Formation overlies unconformably the Upper Eocene carbonates (Ismail and Abdel Ghany, 1999). In the Cairo-Suez province, most of the Miocene sediments have been partially eroded away. By contrast, the Oligocene Gebel Ahmar Formation is often present and locally capped by Early Miocene basaltic ﬂows (Fig. 2A). The Oligocene Gebel Ahmar Formation is well exposed at West of the Gulf of Suez (Fig. 1). Previous studies describing the formation have characterized the deposits a being consists of varicolored, cross-bedded, medium to coarse-grained sands and blackened chert pebbles (Shukri, 1954) with abundant petriﬁed wood (Cuvillier, 1930). 3. Methodology The Oligocene multistorey alluvial complexes in the Cairo-Suez province are well exposed and accessible in several localities

especially along the Kattamiya-Sukhna road. This road exposes transects of 5- to 30-m-high with up to 500 m lateral continuity for the Oligocene succession (Fig. 1C). Perpendicular gullies through these transects provide a three dimensional perspective of the succession. This unique 3D perspective allows the studying the spatial and temporal variability of the Oligocene succession. Field observations of sedimentary facies, bounding surfaces, geometry and architectural characteristics are documented by three proﬁles measured and described at the main road and gullies (Figs. 6–13 ). The proﬁles were described here forming a curved serrated line about 200–900 m in length, oriented mainly east-west (Fig. 1C). The proﬁles are numbered 1, 2, and 3. The proﬁle 1 is located at the western part, whereas the reminder of the proﬁles are located to the eastern part of the study area (Fig. 1C). Figs. 6–8 illustrate proﬁle 1, oriented variably from low to high angle to paleoﬂow. The proﬁle 2 (Fig. 9) is located south of Kattamiya-Sukhna road and is oriented variably from perpendicular to paleocurrent in the east to 20 –30 in the west. Proﬁle 3 (Fig. 10) is located to the north of Kattamiya-Sukhna road being serrated and oriented variably from low angle (20 ) to perpendicular to paleoﬂow. The paleoﬂow directions are parallel to the small gullies of the cliffs in proﬁles 2 and 3. Data was analyzed to establish the depositional

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Fig. 5. Architectural elements of the Oligocene Gebel Ahmar Formation.

paleoenvironment responsible for generating the preserved stratigraphic architecture. Graphic logs and architectural panels record the variety of sedimentary facies present, their building into sets and cosets, the arrangement of such sets/cosets into barforms. The orientation of barform accretion sets were evaluated in relation to the paleocurrent direction to help identify the dominant mode of accretion (e.g., lateral or downstream accretion) within stories. A storey consists of channel incision event with a basal scour surface and comprises the main units of each ﬂuvial complex (cf. Jensen and Pedersen, 2010). Paleocurrent data were measured using mainly planar and trough cross-strata which are the product of in-channel lunate dunes and downstream accretion (DA)

channel bar features. From integrated sedimentary facies, barforms, and bounding relationships, storeys observed in this study were assembled to identify the genetically related ﬂuvial archetypes. Thicknesses of cross-sets were measured to estimate average dune height and bankfull water depth (e.g. Leclair and Bridge, 2001; Reesink and Bridge, 2009; Ashworth et al., 2011; Reesink et al., 2014). Sedimentary facies were deﬁned mainly on the basis of grain-size, sedimentary structures, and contacts. The facies were grouped in facies assemblages, which correspond to a group of facies genetically related to one another, representing sub-environment within a depositional system (Collinson, 1996; Dalrymple, 2010). The sedimentological results are used to deduce

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Fig. 6. (A) Panel sketch for the Oligocene succession at Western part of the study area shows the distribution of Storeys 1–4. The inset is satellite map for western area (proﬁle 1). (B) Photomosaic for the Oligocene Gebel Ahmar Formation at western part of the study area with interpretation of the ﬂuvial architecture. (C) Sketch for the architectures of the Oligocene succession with 2 vertical exaggeration. The orientation of the proﬁle varies somewhat along the cliff, as indicated by the orientation yellow arrows. The bounding surfaces are labeled with numbers in circle. Paleocurrent red arrows indicate the paleoﬂow mean direction from readings within proﬁle. Table 1 General characteristics and interpretation of the sedimentary facies for the Oligocene Gebel Ahmar Formation. Facies

Description

Sedimentary structures

Interpretation

Massive conglomerate Gm

Massive to crudely stratiﬁed pebble to cobble size, clastsupported and poorly sorted (Fig. 4A). The matrix is medium to very coarse sand. Clasts are subangular to well rounded. It is characterized by sharp, well-deﬁned erosional horizontal to inclined contact between beds Clast-supported pebble to cobble conglomerate, beds between 0.2–3 m thick. Clasts are imbricated and stratiﬁcation is well deﬁned (Fig. 4B). Lenses of coarse to very coarse-grained sandstone interbedded. Irregular scours at the base Clast-supported to matrix-supported pebble/cobble conglomerate and coarse to very coarse-grained sandstone alternations with large foreset (>1.5 m, Fig. 4C, D).

Structureless

Longitudinal bars, lag deposits, sieve deposits.

Horizontal stratiﬁcation

Deposited under traction transport by low relief gravel bedforms or as lag deposits. Discontinuous sandstone lenses suggest ﬂuctuations in ﬂow conditions.

Low to highangle-inclined foresets Angle of forests 20–30. Planar crossstratiﬁcation Angle of foresets 15–30.

Migration of gravel bars similar to described by Miall (1977); Rust (1978, 1979); Middleton and Trujillo (1984); Smith (1990)

Trough crossstratiﬁed

3D sandy dunes migrating in the deeper part of channels (Ashley, 1990; Bridge and Gabel, 1992; Skelly et al., 2003).

Trough crossstratiﬁcation Convoluted stratiﬁcation Massive to stratiﬁed

Syn-sedimentary deformation through liquefaction and ﬂuidization caused by high sedimentation rates and sudden decrease in discharge (Owen, 1996; Røe and Hermansen, 2006). Separate set or coset of facies cross stratiﬁed sandstones. Related to bounding surfaces of ﬁrst to third-order.

Stratiﬁed conglomerate Gh

Cross stratiﬁed conglomerategravelly sandstone Gp Planar cross stratiﬁed (pebbly) sandstone Sp Trough crossstratiﬁed (pebbly) sandstone St Convoluted (pebbly) sandstone Std Pebble beds (Facies Gmp)

Medium to coarse-grained, poorly sorted, sandstone. Average set thickness 10–60 cm, maximum set thickness 90–100 cm (Fig. 4B). It is characterized by sharp, horizontal to inclined erosional bases Medium to very coarse-grained, poorly sorted, pebbles are common, sandstone (Fig. 4E). Set thickness 20–100 cm, coset thickness 1 to >5 m. The basal contact is sharp, horizontal to inclined with evidence of erosion between bedsets Coarse to very coarse-grained, sometimes pebbly, poorly sorted. Partly deformed sets of facies Sp and St with gradual transition to strongly deformed structures and massive sand with water escape structures (Fig. 4F). Strings or thin beds of small pebbles. Bed thickness similar to one or a few clasts, which may change laterally (Fig. 4E). Lateral extension from a few metres to more than 100 m. Poorly to well-deﬁned erosional horizontal contact between bedsets

2D sandy dunes & simple bars (Smith, 1970; Cant and Walker, 1978; Ashley, 1990).

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Fig. 7. (A) Storey 1 at the western area (proﬁle 1). (B) The horizontal burrows within ﬁne sandstone bed at the top of the storey. (C) Compound GB element (Storey 1) that consists of stacked unit bars of Gm and Gh facies. (D) Storey 2 at the western area (proﬁle 1 The orientation of the proﬁle varies somewhat along the cliff, as indicated by the orientation yellow arrows. The bounding surfaces are labeled with numbers in circle. Paleocurrent red arrows indicate the paleoﬂow mean direction from readings within proﬁle (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article).

and discuss the possible controlling elements of the valley ﬁlling such as tectonics and climate (e.g. Scherer et al., 2015; Rasmussen, 2014; Almedeij and Diplas, 2005). Facies descriptions and interpretations are presented in Table 1 and Fig. 4. Bounding surfaces were traced and correlated in the ﬁeld. The architectural analysis largely followed the principles of bounding surface hierarchies and architectural elements proposed by Miall (1985, 1988, 1993). Six types of small- to large-scale bounding surfaces, deﬁned as 1–6 orders were identiﬁed. The sixth (6) represents the Oligocene canyon that truncates the underlying Eocene bedrock and it is characterized by strong evidences for subaerial exposure and development of a proper sequence boundary. The ﬁfth (5) order surfaces are laterally extensive erosional surfaces separating individual, large-scale channel ﬁll units. These channel erosional surfaces should ideally be concave-up in strike section but appear horizontal when the channel is larger than the exposure(>200 m). These surfaces can be traced between outcrops, may have local relief and sediment transport directions vary up to 45 across ﬁfthorder surfaces. Fourth-order surfaces represent the upper

boundary of macroforms, and they are typically ﬂat to convexup. Third-order contacts enclose group of elements or complexes, and usually are well-deﬁned planar to slightly curved erosional channels. Second-order surfaces are curved or planar and bound cosets. First-order contacts are planar or curved surfaces that bound individual crossbed set or sets of planebed laminae. 4. Results 4.1. Sedimentary facies Seven sedimentary facies were described and interpreted in the Oligocene Gebel Ahmar Formation (Table 1 and Fig. 4). All of them were dominated by sandstones and conglomerates. Conglomerates (G) are typically clast-supported with a matrix of very coarse to medium quartz sand. Clasts are moderately to well rounded, and imbrication of oblate pebbles and cobbles is common. Over 90% of the clasts were rounded to well-rounded and predominantly consist of chert with some carbonates. Clasts reach

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Fig. 8. (A) Storeys 3 and 4 of the Oligocene Gebel Ahmar Formation at proﬁle 1. Notice the prograding bar macroform of channel C. Close-up view for (B) Channel A and (C) Channel B. (D) Panel sketch for storey 4 at proﬁle 1. (E) Conglomerate-sandstone couplets of storey 4 at proﬁle 1. (F) Large rhizoliths at sandstone interbeds of storey 4, the black mobile pocket (12 cm 5 cm) for scale. The orientation of the proﬁle varies somewhat along the cliff, as indicated by the yellow arrows. Paleocurrent was indicated by red arrows. The bounding surfaces are labeled with numbers in circle (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article).

30 cm in a-axis length, but average 10 cm. The conglomerate normally graded to ungraded. Sometimes, it is interbedded with coarse-grained and pebble sandstones. Three conglomerate lithofacies are recognized; Gm, Gh, and Gp (Fig. 4 and Table 1). The sandstones (S) consist of ﬁne to very coarse-grained textures, sometimes with pebbles and random cobbles. The clasts compositions are the same of those recorded in the conglomeratic facies. Different forms of stratiﬁcation can be used to distinguish four depositional facies; St, Sp, Sm, and Sd (Fig. 4 and Table 1). Trough and planar cross-stratiﬁcations (St and Sp) indicate that these sand bodies were deposited by 3D and 2D migrating large to small-scale dunes, respectively, while massive or structureless bodies (Sm) are likely to be a consequence of rapid deposition (e.g. Todd, 1989; Maizels, 1993; Nichols and Hirst, 1998). Deformed cross-stratiﬁcation (Sd) was formed due to liquefaction. 4.2. Alluvial architectures Nine architectural elements have been identiﬁed within the Oligocene Gebel Ahmar succession (Fig. 5) based on the schemes of Miall (2010a,b, 2014) and Colombera et al. (2012a,b, 2013). These architectural elements include gravel bedforms (GB) such as gravel bars and sheets, gravel-sand foresets, stacked gravel-sand couplets, scour pool ﬁlling, Downstream accretion (DA), Downstream lateral accretion (DLA), lateral accretion (LA), sandy bedforms (SB), Channel (CH), and Hollows (HO) elements. Not all of these features were observed in the same proﬁle. Identiﬁcation and interpretation of the described bounding surfaces are based on the

classiﬁcation of Miall (1996, 2010b). Architectural elements are distinguished by a variety of characteristics including bounding surfaces, lithofacies, geometry, vertical and lateral associations, and paleocurrent orientations. Therefore, storeys are deﬁned as channel-belt that are made up of architectural elements, which in turn are composed of smaller, internal features. Although the 3D nature of the studied proﬁles yields exceptional detail with regard to internal facies, the lateral extent of elements examined remainsrelatively poorly constrained because many of the larger elements exceed the extent of the studied proﬁles; dimensions of these elements are therefore either partial or unlimited (sensu Geehan and Underwood, 1993). Within the following interpretations, approximate paleocurrent velocities have been estimated with reference to grain size and critical erosion velocity analysis ( Sundborg, 1956). Approximate original mean dune height can be calculated using the following equation (Leclair and Bridge, 2001): Hm = 2.9 (0.7) Sm where Hm and Sm represent minimum dune height and mean cross-set thickness, respectively. Since the mean grain sizes associated with architectural elements vary from medium sand to pebble, the minimum paleocurrent velocity is 0.6 m/sec. 4.2.1. Gravel bedforms 4.2.1.1. Gravel bars and sheets (GB). Gravel bar and sheets element comprises massive and stratiﬁed conglomerates (Gh and Gm) facies with subordinate massive and cross-stratiﬁed sandstone (Sp,

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Fig. 9. (A) Photomosaic for the Oligocene Gebel Ahmar Formation at proﬁle 2. inset satellite image for proﬁle 2 shows the location of C-D. (B) the architectural elements of storey 2 at proﬁle 2, shown with x3 vertical exaggeration. (C, D) Sand bedforms SB element of storey 2. (E, F) Trough cross-stratiﬁed sandstone (St) facies lenses with reworked intra-basinal calcrete and extrabasinal chert lags. The orientation of the proﬁle varies somewhat along the cliff, as indicated by the yellow arrows. Paleocurrent was indicated by red arrows. The bounding surfaces are labeled with numbers in circle (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article).

St, and Sm) facies, up to 0.2 m thick (Figs. 5 A, 7C, 11 A). This element consists of lenticular to sheet-like bodies, range in from <0.2 m to about 5 m thick but are typically 0.3–1 m thick. It is locally deformed or convoluted due to syn-sedimentary faulting (Fig. 12B) The width/thickness ratio of the gravel bar and sheets varies from 10:1 to 40:1 within the study area. This element has sharp erosional bases and commonly erode sand bedforms and channels. The erosional relief of the gravel bar and bedforms varies between ﬂat to irregular scour. Some examples of this element at proﬁle 1 have convex-up to ﬂat upper boundaries and show downstream accretions (Fig. 8A). Other examples show no evidence for lateral or downstream accretion pattern (Fig. 7A & B). Sandstone bodies shows lenticular geometry that varies between 1 to 20 cm thick and lateral extension varies between 1 m to 30 m. The sets of the Sp facies show mainly NNW to NW paleocurrent direction. At proﬁle 2, The lenticular sandstone bodies contain reworked calcrete fragments, up to 0.4 m in size (Fig. 9E & F). This element is interpreted to have been deposited as bedload sheets, channel lags and longitudinal bars on a middle to upper or proximal braided plain complex (Rust, 1972, 1978; Smith, 1974; Boothroyd and Ashley, 1975; Boothroyd and Nummedal, 1978; Kleinspehn et al., 1984; Lunt et al., 2004; Long, 2006). The gravel bars include simple unit bars accreted that range between 1–1.2 m heights and 20–50 m length during the ﬂooding stage. Sometimes, the bedload sheets and gravel unit bars are accreted to form

compound bars or prograding downstream as prograding bars. The minimum channel depth can be estimated based on the preserved unit bar thickness between 1.1 to 3 m (e.g. Lunt et al., 2004). In the Sagavanirktok river (N Alaska), compound bars are hundreds of meters in width and length and have heights of 1–3.5 m although the channel width about 3.9 m. This gravel bars and sheets are accumulated during the ﬂooding stage. During the waning stage, the sand accumulated on the gravel bars as bar top. The presence of reworked calcrete fragments reﬂects the reworking of paleosols and the formation of channel-ﬁll successions that include both reworked calcrete and extrabasinal clasts (Figs. 9E & F). During ﬂood events, these channels have a high erosion capacity, and can erode calcretes that formed in the inactive parts of channel-belt. These deposits are similar to those described by Allen and Williams (1979); Marriott and Wright (1993) in the Old Red Sandstone of Wales; Gómez-Gras and Alonso-Zarza (2003) in Permian and Triassic deposits of Balearic Islands (Spain). 4.2.1.2. Gravel-sand foresets (GSF). Example of this element is described at proﬁle 2 and typically consists of large-scale inclined gravel and sandstone foresets (Figs. 4 C, 5 B). The sandstone set is mainly composed of coarse-grained cross-stratiﬁed Sp facies, 0.2– 0.8 m thick, while the conglomerate bodies consists of Gm and Gh facies, 0.2–0.5 m thick. The conglomerate bodies are mostly convoluted and deformed especially to the south (Fig. 10E & D).

S.S. Selim / Proceedings of the Geologists’ Association 128 (2017) 234–255 Fig. 10. (A) the alluvial architectures of Oligocene succession at proﬁle 3, inset satellite image for proﬁle 3 shows the location of B–E. (B) Panoramic view for the eastern part of the cliff (right rectangle of Fig. 5A). (C) Panoramic view for the western part of the cliff (left rectangle of Fig. 5A). (D, E) The architectures of Oligocene at proﬁle 3 (orientation N330-N320). The orientation of the proﬁle varies somewhat along the cliff, as indicated by the yellow arrows. Paleocurrent was indicated by red arrows. The bounding surfaces are labeled with numbers in circle (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article).

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Fig. 11. The different architectural elements of Storey 1 at proﬁle 3: (A) The Gravel bars GB and DLA elements of storey 1. (B) 3D prespective of DLA element. (C) Downstream DA element of alternate or transverse bar scoured with laterally superposed small scale Hollow elements (HO). (D) Close up view for the Hollow element shows the basal fourth order bounding surface covered with one pebble thick string (black arrows). Notice the paleocurrent of the HO element is almost against the paleocurrent reading from the DA element. The orientation of the proﬁle varies somewhat along the cliff, as indicated by the yellow arrows. Paleocurrent was indicated by red arrows. The bounding surfaces are labeled with numbers in circle (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article).

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Fig. 12. (A) Downstream accretion DA element of storey 2 (proﬁle 3). (B) The vertical attitude of the storey 2 elements due to NW-syn-sedimentary faulting.

The height of the foresets is commonly 2.5 m, but heights of up to 5 m occur (Fig. 10). The lateral extent of the foresets is in the order of 2 to a 10 m, and the foresets dip in various directions (W, NW, N; Fig. 10). The bedding surfaces dip 20 –25 , but pass laterally into progressively more gently dipping sets of master bedding, which typically dip at 10 –15 , but rarely at less than 5 (Fig. 12). This element rests on concave up, basal deep erosional, ﬁfth-order surface (Fig. 10) that causes completely missing of the underlying elements (i.e. DA1 and SB elements). This element is interpreted as large-scale mid-channel bar deposits in deep gravelly channel (e.g. Bristow 1987; Ashworth et al., 2000; Best et al., 2003; Sambrook Smith et al., 2006; Smith et al., 2009; Allen et al., 2013). Large cross-stratiﬁcation dips commonly up to the angle of repose reﬂects the downstream bar migration of deep channel due to avalanching down the slipface of the bar, causing bar migration. The accretion toward NW, W, and N directions reﬂects the downstream to slightly oblique to the paleoﬂow direction, as is commonly seen in mid-channel bars (Sambrook Smith et al., 2006). Similarly, Smith et al. (2009) described large scale forests in the largest braided river such as Río Paraná, Argentina as prograding or accreting bars. Based on the preserved bar thickness, the expected minimum channel depth 5–10 m (e.g. Lunt et al., 2004). The deformed and convoluted conglomerate bodies liquefaction due to syn-

depositional faulting, that affected the underlying elements (see Fig. 12B). Similar elements were described in other braided rivers such as Jamuna (Best et al., 2003), Wisconsin (Mumpy et al., 2007) and South Saskatchewan (Sambrook Smith et al., 2006) rivers. Alternatively, Wooldridge and Hickin (2005) described downstream steeply inclined gravelly bedforms as delta foresets in gravelly Fraser and Squamish rivers (British Columbia, Canada). 4.2.1.3. Gravel-sand couplets (GSC). The examples of this element is widely distributed in the study area, 2–5 m thick and can be traced laterally for 2 km wide between isolated hills. It entirely consists of stacked conglomerate-sandstone couplets (Figs. 5 C and 13 A). It is characterized by shallow, erosive, ﬁfth-order basal surface. The conglomerate sheets consist of erosively to ﬂat-based, continuous and discontinuous bodies of Gm and Gh facies, up to 0.7 m thick and >100 m lateral extension (Fig. 13B). The sandstone sheets consist of cross-stratiﬁed sandstone (St and Sp facies), decimeters to meters thick. Locally, the sand sheets are characterized by large rhizoliths ranging between 20 to 100 cm length and 2 to 5 cm diameter (Fig. 7F). The suite of characters shown by this element suggest that the gravel-sand couplets resulted from sheet to broad channel ﬂow processes. This element is similar to conglomerate-sandstone couplet of the Donjeck model of Miall (1977, 1978); Smith (1990).

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Fig. 13. (A) Stacked sandstone-conglomerate couplets of storey 4 at proﬁle 3. (B) Storey 4 rests directly on the Upper Eocene.

According to this model, each couplet was deposited in a broad, poorly conﬁned stream in which gravel accreted in the lowest parts of the channel under high energy ﬂows. This followed by waning of the ﬂow and sand deposition in the topographically higher parts. Similar couplets have been described by Blair and McPherson (1994), who interpret the couplets as due to changes in ﬂow hydraulics, initiated by ﬂow expansion and decrease in slope gradient. Similarly, the couplets also suggest low turbulence in the ﬂows that enabled the stepwise deposition of gravel and sand with fall in the ﬂow capacity (see Chamyal et al., 1997). The presence of large rhizoliths at the top of some sandstone sheets reﬂects the dominance of vegetation on the inactive parts of the alluvial system. 4.2.1.4. Scour pool ﬁlling (SPF). Examples of this element are common at proﬁle 3 and typically consists of a symmetrical or asymmetrical scour or a number of overlapping scours (Figs. 5 D, 10 D). The geometries of this element are long and deep shapes of cutand-ﬁll. They are characterized by an erosive, concave-up lower bounding surface and a usually roughly horizontal upper surface (Fig. 10D & E). The internal structure of this element consists of ribbon geometry of Gm and Gh facies with only minor occurrence of stratiﬁed and massive sandstone (St/Sm) facies. This element is interpreted as scour pool ﬁlls formed at channel conﬂuences or channel bends (e.g. Ashmore, 1982). The massive inﬁlls are most likely to represent basal channel ﬁll deposits (Marren, 2001). The forms of scour features depend on channel geometry, conﬂuence angle, discharge, sediment load, and lateral or longitudinal shifts of the scours (see Best, 1988; Ashmore, 1993; Siegenthaler and Huggenberger, 1993). This element is common in braided rivers (Best 1988; Best and Ashworth 1997). Similarly, Kostic and Aigner (2007) described chute channel or channel conﬂuences of gravelly river in Neckar valley (SW Germany) using the Quaternary outcrops and GPR characteristics.

4.2.2. Sand bedforms 4.2.2.1. Downstream accretion (DA). This element is the most common architectural element in the study area and it is including two types; simple downstream accretion DA1 and complex downstream accretion DA2 (Fig. 5E). i. Simple downstream accretion (DA1). The example of this element is described at proﬁle 3 and it consists mainly of solitary set of Sp facies (Figs. 5 E & 11 C). This set is composed of large-scale planar cross forests, up to 2 m thick. Foresets are inclined at angles of up to 25 , whereas set bounding surfaces are typically horizontal to inclined at low angle up to 10 . It rests on ﬁfth to sixth order bounding surface. This element is incised by at least two hollow elements (HO, Fig. 11C & D). The large-scale (2m) steeply dipping foresets were formed by down accretion well-developed slipface of possible migrating alternate bar or transverse bar within a comparatively large and deep ﬂuvial channel (c.f. McCabe, 1977; Collinson, 1996; Collinson et al., 2006; Miall, 2010b). The uniform paleoﬂow direction towards the NW with the paleoﬂow, indicate bar migration parallel to the main ﬂow direction. This element is probably interpreted as bank attached bar due to the uniform dip direction. Bar height may vary between half and bankfull depth (cf. Bristow, 1987; Bridge, 2003; Reesink et al., 2014). Therefore, the preserved planar cross-bedded bar thickness (ca 2.0 m) suggests that the depth of the host channel was in the area of 2 m–4 m (cf. Bristow, 1987; Bridge, 2003; Reesink et al., 2014). The alternate bars may have developed attached to channel bank (e.g. McCabe, 1977; Collinson, 1996; Collinson et al., 2006; Reesink et al., 2014). Alternatively, the alternate barforms may have develop as large mid-channel bars (Collinson, 1996; Collinson et al., 2006; Rasmussen, 2014) or

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may develop as they migrate into a scour pool associated with a channel conﬂuence (McCabe, 1977; Collinson, 1996; Collinson et al., 2006). ii. Complex downstream accretion (DA2). The examples of complex downstream accretion type DA2 represent the most common architectural element in the study area (Figs. 7 D and 9). The most examples of this element up to 5 m thick and consists of cosets of facies St with some 0.05– 0.15 m thick diffusive gravel layers (Gmp) separated by 2nd- to 3rd-order reactivation surfaces (Fig. 7D). Subordinate sets of Sp facies up to 0.5 m thick and >50 m wide are described at proﬁle 1 (Fig. 4B). Low-to high-angle inclined foresets that form up to 1 m thick and >1.5 m wide trough cross-bed sets that overlapped vertically and laterally to form compound cosets of sandbodies up to 5 m thick and >100 m wide (Fig. 7D). At proﬁle 2, examples of this element shows upward decrease in the crossbed set thickness. The examples of this element at proﬁle 1 are characterized by dominance of vertical and inclined traces such as Skolithos ichnogenera especially along the 2nd and 3rd-order reactivation surfaces (Fig. 4B). Second and third-order accretion surfaces between superimposed bedforms dip towards N320–340. The mean sediment transport direction is measured NW (Figs. 7 D and 9 C & D). The DA2 element shows good evidence for downstream accretion as the foresets are inclined at 15–25 and the set bounding surfaces are inclined in the NW and NNW trends. The internal arrangement of facies within examples of this element records trough cross-bedded sets that are indicative of a generally northerly migration of medium to large-scale sinuouscrested dunes. Sandbodies accumulated and accreted downstream on a channel-ﬂoor setting or on gravel bars/sheets (GB) elements, this interpretation is similar to features described by Bristow (Best et al., 2003) who envisaged downstream accreting mid channel bars that developed in the Jamuna River (Bangladesh) and the bedforms that described from the Lower Devonian Brownstones (Welsh Borders, Allen, 1983). The average preserved set thickness 0.5 m, for facies St suggests minimum mean dune heights of 1.1 m– 1.8 m (Leclair and Bridge, 2001). Cant and Walker (1978) indicates that the dunes with a maximum height of 1–5 m developed during ﬂood events within channels possessing a mean ﬂow depth of 3.0 m, whereas deeper channels possessed larger dunes (cf. Reesink and Bridge, 2009; Ashworth et al., 2011; Reesink et al., 2014). The upward decrease in set thickness reﬂects decrease in channel depth through time (c.f. Reesink and Bridge, 2009; Ashworth et al., 2011; Reesink et al., 2014). Although experimental studies of Coleman (1969) suggest that dune height may be less than or equal to the water depth, ﬂume experiments by Reesink and Bridge (2009) deduced that the height and length of sand dunes increase with water depth and discharge rate. The macroform probably developed around a simple bedform nucleus as water levels rose (Cant and Walker, 1978; Alexander et al., 1994; Bridge et al., 1995, 1998). The occurrence of compound sets of St and Sp facies as well as thin gravel layers within DA2 suggests that the variable discharge during deposition. The presence of vertical and inclined traces such as Skolithos reﬂects the energy ﬂuctuations during the channel deposition. Worm response to water table ﬂuctuation led to formation of traces (Fitzgerald and Barrett, 1986). 4.2.2.2. Downstream and oblique accretion (DLA). The example of this element were described at proﬁle 3 and it consists of packages of cosets of St, Sd, and Sp facies being separated by 1st to 3rd-order bounding surfaces to the east of the proﬁle 3 (Figs. 5 F and 11 A). The average St set thickness is up to 0.8 m and width >50 m. The ﬁrst and second-order bounding surfaces dip in northerly direction. Second-order accretion surfaces between small-scale superimposed bedforms, as well as the third-order reactivation surfaces dip in

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westerly directions (Fig. 11B). The third-order reactivation surfaces are traceable across the cliffs and perpendicular gully (Fig. 11B). Sediment transport directions are towards NE. Locally a one-clast thick bed of pebbles separates sets of facies St and Sp. The internal facies within the example of this element records trough and planar cross-bedded sets that are indicative of a generally northerly migration of medium and small-scale straight and sinuous-crested dunes. These sandbodies accumulated in a channel-ﬂoor setting above gravel bar (GB) element. The second and third-order bounding accretion surfaces reﬂect that the dune barforms are accreted westerly (i.e. obliquely to main paleoﬂow). The gravel lags indicate the ﬂuctuation of the ﬂow energy possibly with increase in paleo-discharge and an increased rate of dune migration and accumulation (Ashworth et al., 2011). These dunes were developed in channels with ﬂow depth around 3 m (cf. Cant and Walker, 1978; Reesink and Bridge, 2009; Ashworth et al., 2011). The preserved set thickness is 0.8 m, for St facies, suggests original mean dune heights of 1.7–2.8 m (Leclair and Bridge, 2001). The abundance of deformed stratiﬁcation of Sd facies is due to liquefaction, and syn-depositional faulting. 4.2.2.3. Lateral accretion (LA). The example of this element is described at proﬁle 2 (Figs. 5 G, 9 A & B) and it laterally interﬁnger with the DA2 bedforms. The LA element consists of alternated large-scale inclined coarse-grained cross-stratiﬁed to massive sandstone with gravel layers of about 3 m thick, dipping about 15 , perpendiculars to paleoﬂow direction (Fig. 9B). This element represents the lateral accretion of in-channel bar (e.g. Lunt et al., 2004; Wooldridge and Hickin, 2005). The inclination of the bedding reﬂects accretion due to bar migration. the lateral accretion pattern represents the side growth of compound bars that described in many recent analogues of large sandy and gravelly braided rivers such as Sagavanirktok (northern Alaska, e.g. Lunt et al., 2004; Sambrook Smith et al., 2006); Jamuna (Bangladesh, Best et al., 2003); and Río Paraná, Argentina (Smith et al., 2009), Wisconsin (Mumpy et al., 2007), Brahmaputra River (Bristow, 1993). Lateral accretion is locally signiﬁcant in the enlargement of mid-channel bars in braiding and wandering rivers (Smith, 1974; Ori, 1982; Ramos and Sopeña, 1983; Ashmore, 1991). It is a signiﬁcant and widespread depositional style in braiding rivers with slowly migrating bars (Lunt et al., 2004). The accretion of sediment onto gently dipping bar-tail margins is important in directing down-bar growth during some phases of bar migration (Smith, 1990; Lunt et al., 2004). Alternatively, this element is considered as bars which formed in a locally more sinuous channel tract (L. Devonian Brownstones, welsh borders, Smith, 1990). 4.2.2.4. Sandy bedform (SB). The examples of this element are described at proﬁles 1 and 3 (Figs. 5 H, 7 A–C). and it consists of mainly of planar cross-stratiﬁed sandstone Sp facies with subordinate gravel trains. The set attains 0.1–0.5 m thick and the cosets varies between 0.5–1.2 m thick, with internal reactivation or bounding surfaces of ﬁrst to 2nd order, dip easterly and westerly. These sandstone bodies have width greater than the outcrop length (>100 m, Figs. 6 and 13). The third order surfaces are associated with gravel trains or sheets vary from one to several clasts thick. This element shows sharp ﬂat or erosive bases. The examples of this element mostly rest on gravel bedforms (GB). At proﬁle 1, the example of this element is characterized by Treptichnus horizontal cylindrical ichnogenera at the top (Fig. 7C). The lithofacies that comprise this element are indicative of the migration and accumulation of small to medium-scale straightcrested dunes accreted on gravel bar (GB) top easterly and westerly. The dominance of Treptichnus horizontal ichnogenera reﬂects the channel deactivation and prevailing of the quite conditions (cf. Buatois and Mángano, 2004; Buatois et al., 2002).

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Table 2 The description and interpretations of the Oligocene incised valley ﬁll storeys at the western area. Storey

Architectural elements

Storey 1

GB Subordinate SB DA2 subordinate GB Multistorey CH

Storey 2 Storey 3 Storey 4

GB

Dimensions

Sedimentary facies types

Biogenic and chemical sedimentary structures

Interpretation

Width

Thickness

>100 m >50 m

3m 1m

Gm, Gh, and subordinate Sp/Sm

Treptichnus ichnogenera

Gravely braided streams

>200 m >50 m

5m 0.5–1 m

St, Sp, Sd, with subordinate Gmp, Gm

Skolithos ichnogenera

Sandy braid plain

50– 100 m >200 m

2–3.5

St, Sp, GB

Calcrete

Sand-dominated channel ﬁll

6m

Gh, Gm with subordinate Sp and Sm

Large rhizoliths wood fragments (Cuviller, 1930)

Stacked prograding bars of gravely shallow braided channel

Table 3 The description and interpretations of the Oligocene incised valley ﬁll storeys at eastern area. Storeys

Architectural elements Dimensions

Storey 1 DLA DA1 SB Subordinate GB OH Storey 2 DA2, LA, subordinate GB Storey 3 GSF subordinate SPF Storey 4 GSC

Sedimentary facies

Biogenic and chemical sedimentary structures Interpretation

3–4 m 1.5–2 m 1–1.2 m 0.5–1.5 m 0.3 m

Sp, St, Sd with subordinate Gm/Gh

Skolithos ichnogenera

Sandy braid plain

5m 2–3 m 1m 1–5 m 1–1.5 m

St, Sd, Sp, Gp, with subordinate Gm

Skolithos ichnogenera Reworked calcrete

Sandy braid plain

Width

Thickness

>100 >100 m >100 m <50 10– 15 m >200 m >50 m >50 m >100 m >15 m >500 m

>5 m

gravelly braided channel ﬁll

Gp, Gm, Gh,

Gm, Gh, Sp, St

Large rhizoliths

Gravelly unconﬁned braid plain (probably on top of alluvial fan)

4.2.2.5. Channel (CH). Two examples of this element are described at proﬁle 1; CH A and CH B (Figs. 5 I, 8 A & B). Externally, it is characterized by concave-up channel to wedge shaped geometries of up to 2–3.5 m and <100 m width. CH element usually forming a multi-storey channel geometry (Figs. 8A & B). CH A element consists of cosets of planar and tabular cross-stratiﬁed sandstones (0.2–0.8 m) that characterized by westerly dipping. The sets are bounded by 1st/2nd bounding surfaces of westerly and Northerly dipping. CH B consists of large broad shallow concentric concaveup sets of Sp and Gm facies (Fig. 8C). CH element have sharp erosional bases with relief of about 0.5–3 m and erodes the underlying sand bedforms and other channels. The top of CH B is characterized by white calcareous nodules and patches from millimeter to centimeter scale (calcrete, Fig. 8C). This element is interpreted as a channel ﬁll. Channels comprising facies Sp records deposition. The internal facies within examples of CH A element records medium and small-scale planar cross-bedded sets that are indicative of a generally westerly migration of small and medium-scale straight-crested dunes. These dunes were accumulated in a channel-ﬂoor setting that itself formed a fourth to ﬁfth-order bounding surface (the base of the element). The concave-up stratiﬁcation of CH B reﬂects the deposition from suspension during waning ﬂoods in chute channel (c.f. Kostic and Aigner, 2007). The channel geometry probably represents channels with nearly N/NW paleocurrent directions; the E–W orientations of the exposures display the channels in almost transverse section. The patches of calcrete reﬂects the formation of transported and re-cemented calcrete.

ribbon-like geometry of up to 0.5 m thick and 10 m width. Internally, it consists of single set of cross-stratiﬁed sandstone (Sp) facies. The basal boundary is covered by gravel lags. This element is interpreted as hollow element (e.g. Miall and Jones, 2003) or chute channel ﬁll. HO element represents the rapid cut and ﬁll of scour hollows at bars or channel conﬂuences downstream from bars or bars (Miall and Jones, 2003). The occurrence of erosional scours, planar cross stratiﬁcation and their concave-upward lower bounding surfaces of individual channels are evidence of deposition by vertical aggradation (Campbell, 1976; Collinson, 1978). Similar element was demonstrated by Singh and Bhardwaj (1991) at Ganga river, India. Various architectural elements form together channel-belt or storey that is bounded by a laterally extensive, erosive basal surface of 5th to 6th order. The spatial arrangement of elements within the storey is the result of the spatial relationship between major morphologic units and the movement of channels and bars within a channel-belt (Miall, 1994). The major bounding surfaces deﬁne the tops and bases of the individual storeys or primary channels. Four storeys (1–4) are tentatively identiﬁed in the western area (proﬁle 1, Figs. 6–8 and Table 2) and in the eastern area proﬁles 2 and 3 (Figs. 9–13, Table 3). Generally, the facies of Oligocene deposits at proﬁle 1 are characterized by ﬁner grain size and trace fossil abundance.

4.2.2.6. Hollow (OH). This element is uncommon in the study area and described at proﬁle 3 (Figs. 5 J, 11 C & D). It is characterized by

Two main facies assemblages (FA) can be recognized in the Oligocene Gebel Ahmar Formation.

4.3. Facies assemblages

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The FA1 dominates the studied proﬁles, accounting for 70% of the exposed section, especially in storeys 1 and 2 (Figs. 6, 9–11). It is composed of sandstone-dominated (ca 80%) packages with subordinate conglomerate (ca 20%) have maximum observed thickness of 10 m. Sandstone bodies consists of planar and trough cross-stratiﬁed sandstones (Sp and St facies). Conglomerate bodies consist of lenticular and sheets of massive and stratiﬁed conglomerates (Gm, Gmp and Gh facies). It is more abundant in the lowermost part of this assemblage, and upward ﬁning is observed through this assemblage. The architectural elements dominate this assemblage includes sand bedforms such as downstream accretion (DA1, DA2, Figs. 9 C–D,11 C, 12 A), lateral accretion (LA, Fig. 9A and B), downstream-oblique accretion (DLA, Fig. 11A & B), sandy bedform (SB, Fig. 11B & C), and Channel (CH, Fig. 8A–C), gravel bars and sheets (GB, Figs. 9 C & D,11 A, 12 A) with subordinate hollow (HO, Fig. 11C & D). Measurements of trough cross-strata axes indicate an overall paleocurrent direction to the north and northwest with local paleocurrent directions toward the northeast. This assemblage is interpreted as sandy braid plain deposits (e.g. Boothroyd and Ashley, 1975; Miall, 1977) because constituent architectural elements are indicative of various within-channel elements (e.g. Soltan and Mountney, 2016; Ashworth et al., 2011). The ﬁning upward pattern is interpreted as representing an individual stream which comprises channels and bars (cf. Jo et al., 1997). The interbedded pebbly trough-cross bedded sandstone associated with the channel-ﬁll conglomerates indicates discharge ﬂuctuations of the stream (e.g. Steel and Thompson, 1983). Thus, the entire association suggests prolonged stream energy weakening (e.g., Bridge and Gabel, 1992; Skelly et al., 2003; Jensen and Pedersen, 2010). Within individual channel-belt deposits, assemblages of downstream, lateral, and Downstream-oblique elements mainly reﬂect lateral superposition of different types of bars and channels. The constituents of storeys 1 and 2 reﬂect the channel belt is characterized by both stacked and laterally migrated broad channels. The ﬁll of each channel belt or storey begins with ﬂooding event that associated with accumulation of gravel bars and sheets (GB). The conglomerate in the basal part of ﬁning upward assemblage represents the coarsest particles transported by stream ﬂows, and is comparable to diffuse gravel sheet (Hein and Walker, 1977; Lunt et al., 2004) or channel lag sheet (Nemec and Postma, 1993), longitudinal bars that lacking well-developed slipfaces (cf. Boothroyd and Ashley, 1975; Hein and Walker, 1977). It was settled in frictional freezing of the gravel traction carpets during waning of the stream ﬂood. This followed by waning of ﬂoods and accumulation of large to medium scale sand dunes of DA1, DA2, SB, DLA or a combination. The development of compound bars by unit bar accretion or coalescence is a common process in braided rivers (e.g., Cant and Walker, 1978; Bristow, 1987; Sambrook Smith et al., 2006). The trough cross-stratiﬁcation can be produced from various bar forms such as mid channel bars, while the planar cross-stratiﬁcation can be produced from various bar forms such as mid channel, alternate, and bank attached bars (e.g. Smith, 1971, 1972; Boothroyd and Ashley, 1975; Todd, 1996). The dominance of the large-scale bedforms at storey 1 and storey 2 of the eastern area reﬂects the deposition in channels with depth 2–4 m, the upward decrease in the dune thickness reﬂects the upward decrease in the channel depth (c.f. Reesink and Bridge, 2009; Ashworth et al., 2011; Reesink et al., 2014). Predominantly Northwest/North northwest to north paleocurrent indicate northerly migrated large- to small-scale dunes within almost consistent paleoﬂow and channel trend. FA 2 represents storeys 3 and 4 and forms up to 30% of the overall section (Figs. 6 B & C, 7 E and 14 ). It comprises conglomerates facies (Gm, Gh, and Gp) with subordinate trough and planar cross-stratiﬁed sandstone facies (Sp, St, Sm), Fig. 4 and Table 1. The assemblage shows general ﬁning upward pattern as

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seen in FA 1. Conglomerate-dominated (ca 50–90%) packages with subordinate sandstones (ca 10–50%) have maximum observed thickness of 6 m. Conglomerate-dominated bedforms are dominated by gravel bars and sheets (GB, Figs. 7 & C, 8 A & E), gravelsand foresets (Figs. 4 C, and 10 C–E), gravel-sand couplets (Fig. 13), with subordinate scour pool ﬁlling architectural elements (Fig. 10D & E). The conglomerate bedforms are commonly sheet-like bodies and bounded by sharp bases (Fig. 7E). the gravel-sand foresets are commonly wedge-shaped, 1–5 m thick and up to 50 m in lateral extent (Fig. 10C–E). GB element is the most common gravel bedforms of the Oligocene Gebel Ahmar Formation. Individual GB are generally 0.2–1.5 m thick and extend laterally for tens of meters; thicker beds, up to 5 m are commonly amalgamated. The sandstone includes lenticular and sheet-like bodies, 0.1 m–1 m

Fig. 14. Depositional model for the Oligocene Gebel Ahmar Formation.

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thick. The lower boundary is either gradational to the underlying conglomerate or sharp (Fig. 7E). FA2 is dominated by thin SB element with common horizontal Treptichnus ichnogenera (Fig. 7C). Common large rhizoliths are encountered in the gravel-sand couplets of storey 4 (Fig. 8F). The FA 2 shows characteristics of gravelly braided stream deposits based on the channel-belt deposits are mainly gravels with subordinate sands and the prevailing of the in-channel bedforms such as gravel bars and sheets, gravel-sand foresets, and gravel-sand couplets. The sheet-like geometry (>100 m) is indicative of gravel sheets or low relief longitudinal bars (Boothroyd and Ashley, 1975; Todd, 1989; Brierley et al., 1993; Nemec and Postma, 1993; Jo et al., 1997) and/or thin bedload sheets (Reid and Frostick, 1987). The poor development of such gravels bar may be responsible for lacking of the cross-stratiﬁcation, that reﬂects shallow depth with respect to grain size. The gravel bedforms are formed when the ﬂows on the gravel threshold (e.g. Lunt et al., 2004). The channel ﬁll composed of unit bars, that coalesced to form compound bars. During the falling and low-ﬂood stage, the sands was mostly deposited (e.g. Jo et al., 1997; Nichols and Hirst, 1998; Lunt et al., 2004). Sandy small-scale cross-sets mostly occur as overbank deposits capping channel or form drapes in the troughs of unit bars (Lunt et al., 2004). The presence of the largescale gravel-sand forests, up to 5 m thick, of storey 3 reﬂects the progradation of large-scale gravelly mid channel bars. These bars probably developed in wide and deeper channels (>5 m depth). The large mid channel bars were cut by many chute channels with development of SPF element. The syn-sedimentary active fault movements can also cause both deepening and/or widening of valley (e.g. Ardies et al., 2002). This was evidenced by synsedimentary faults (Fig. 12B) and deformed gravel foresets (Fig. 10D & E). The gravel-sand couplets of storey 4 with up to 2 km lateral extension reﬂects shallow broad gravelly channels based on wide extension barforms, and lithofacies constituents. Streams with high energy can transport a mixture of gravels and sands sheets (Reid and Frostick, 1987). The FA 2 similar to gravelbed streams probably in alluvial fan and outwash fans (e.g. McGowen and Groat, 1971; Boothroyd and Ashley, 1975; Allen, 1981; Ori, 1982; Todd, 1989; Evans, 1991; Brierley et al., 1993; Maizels, 1993; Jo et al., 1997). 5. Discussion 5.1. Paleogeography and depositional model During the Eocene, northern Egypt was characterized by passive margin with prevailing the carbonate deposition (Dolson et al., 2005). By the early Oligocene, a major eustatic sea level fall due to the climatic cooling (Miller et al., 2008). The sea level fall was subsequently followed heavily karstiﬁcation and incision of the Eocene carbonates (Dolson et al., 2005). An Oligocene incised valley-ﬁll was produced through the development of a ﬂuvial system in Cairo-Suez province (Fig. 14). Three recognized phases are: (1) downcutting, (2) ﬁlling (3) expansion (Fig. 14). Downcutting and net erosion of the underlying Middle to Upper Eocene bedrock (shallow carbonate shelf deposits) creating an incised valley resulted from a base-level drop, related to glaciation. The paleovalley was then inﬁlled by widespread braid plains during the subsequent phases of ﬁlling (1–3 storeys) and expansion (Storey 4) phases. Storeys 1–2 represents mainly deep sandy dominated braid plain as conﬁrmed by the prevalence of large downstream, downstream/oblique, and lateral accretion architectural elements throughout the studied area. Storey 1 is dominated by in-channel sandy bedforms such downstream accretion (DA1, DA2, Figs. 5 F, 7 D, 9 C & D, and 11 C), downstream/oblique accretion (DLA, Fig. 11B),

and lateral accretion (LA, Figs. 5 H and 9 A & B), formed by the migration of large scale macroform such as mid-channel, alternate or transverse, and bank attached bars in addition to secondary channel (CH, Fig. 8A–C) and cross bar channel (OH element). These sandy bedforms accreted on and around gravel bars and sheets (GB, Figs. 5 A, 7 C, 8 E, 9, 11 A). Storey 2 is dominated by migration of medium to small scale similar bedforms. Storeys 3 reﬂect the dominance of deep gravelly-dominated braided streams with accretion of giant mid channel gravel bars that transected by chute channels in the eastern area and individual channels in the western area. The channels became deeper and wider due to the synsedimentary faults. Storey 4 (expansion phase) includes sheet-like, poorly conﬁned gravely streams, probably of alluvial fan outwash. The Storeys 1–4 show progressive upward shifting from sandy braid plains to gravelly braid plains and poorly conﬁned gravelly channels probably on alluvial fan surface (i.e. distal to proximal shift in facies). In the braided ﬂuvial systems and the similarities between in-channel bar and ﬁll successions may render such deposits indistinguishable from one another (see Ashworth et al., 2011). The ﬂuvial style reﬂects the interplay between gravelly and sandy braided streams that are composed of sheet-like geometry corresponding to an erosion-dominated, multistorey channel body (as deﬁned by Gibling, 2006). Examples of modern and ancient braided ﬂuvial systems (Clarke and Dixon, 1981; Abdullatif, 1989; Bristow, 1993; Singh and Bhardwaj, 1991; Huggenberger, 1993; Siegenthaler and Huggenberger, 1993; Mosselman et al., 1995; Liu et al., 1996; Anderson et al., 1999; Sambrook Smith et al., 2005; Jones et al., 2001; Skelly et al., 2003; Lunt et al., 2004; Rasmussen, 2014; Soltan and Mountney, 2016, and references therein) demonstrate that such systems are dominated by coarse-grained sediments ranging from sand to gravel grades (Fig. 11A–E). The lacking of the ﬁne-grained deposits is common in these systems such as Oligocene succession of this study, Lower Brimham Grit (Soltan and Mountney, 2016), and modern Brahmaputra (Bristow, 1993) and Gash rivers (Abdullatif, 1989). The similarities between in-channel bar and ﬁll successions across the braid plain led to difﬁculty in the recognition of discrete channel-belts, and only multiple lateral channels ﬁlls and bar deposits could be identiﬁed. In the same context, Miall and Arush (2001) indicated that most or all braidplain systems were accommodated by channels and therefore little room was available for ﬂoodplain accumulation. These observations could indicate that the Oligocene succession represent a braid plain with multiple laterally active channels (Fig. 14A–C). However, most modern braidplain examples have relatively narrow active channel-belts, 1–2 km in width (Bridge and Gabel, 1992; Gabel, 1993; Cole, 1996; Bridge et al., 1998; Skelly et al., 2003; Lunt et al., 2004; Lunt and Bridge, 2004; Sambrook Smith et al., 2006) but may reach up to 10–20 km as in the Brahmaputra River (Thorne et al., 1993; Mosselman et al., 1995; Richardson et al., 1996; Best and Ashworth, 1997; McLelland et al., 1999; Ashworth et al., 2000; Best et al., 2003; Richardson and Thorne, 1998, 2001; Sambrook Smith et al., 2005; Smith et al., 2009; Lahiri and Sinha, 2012; Sarker et al., 2014; and references therein). Paleocurrent data display an azimuthal range from N290 to N80, but the main orientation ranged between N330 and N00 (Fig. 15). At the studied localities the more pronounced paleocurrents are northward. Westward orientation was measured as N320 and N10 and the eastward trends were also between N30 and N60. The paleocurrents are consistent through the Oligocene valley ﬁlling laterally and vertically and show no signiﬁcant variation between storeys. In the study area, the braidplain width cannot be deﬁned exactly due to limited outcropping. However, using the mean ﬂow direction at the eastern and western areas and the areal distribution of the Oligocene sediments, a minimum estimate of braid plain width can be determined. The perpendicular section

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Fig. 15. Paleoﬂow distribution measured at various proﬁles and reported on the satellite image of Fig. 1.

corresponds to a minimum braidplain width of 5 km (Fig. 15). If the braid plain records the lateral amalgamation of multistory channel belts, and because individual channel belts cannot be recognized, the width of each channel belt cannot be assessed accurately. However, the sedimentary succession of the western area is characterized by ﬁner grain size relative to the eastern area and displays abundant vertical and inclined ichnofossils (Fig. 4B). This probably suggests that the western area represents the edge of the channel-belt, thus being signiﬁcantly thinner and ﬁner-grained as sediments were deposited in smaller channels on the periphery of the main active channel-belt. The provenance of the Oligocene valley in the Cairo-Suez province has been inferred from the clast types present in the conglomerates and northerly paleocurrent data (Fig. 14). This suggests that the sediments could have been sourced upstream from the exposed nearby Eocene and the Upper Cretaceous carbonates and Jurassic/Lower Cretaceous clastics at the inverted Galala-Araba structures to the south. This agrees with the conclusion of Said (1981) and Issawi and McCauley (1993). They concluded that the Oligocene clastics are sourced from the Red Sea Hills as the sediments linked with the Red Sea rifting (Fig. 1A & B). Overall it is clear that the source area was located to the south of the basin, but more investigation is needed to fully constrain the source terrane and the dimensions of the drainage basin. 5.2. The interaction between tectonics and sediment supply The incision of the valley described here commenced during climate cooling that was coincident with the glaciation in Antarctica (Miller et al., 2008). A eustatic sea-level fall was thus

important in the timing of the incision. However, the Upper Cretaceous-Paleogene tectonic inversion of hinterland (Hussein and Abd-Allah, 2001; Dolson et al., 2005) as well as the Late Eocene-Oligocene crustal updoming in the Ethiopia (Hofmann et al., 1997) play a major role in the ﬁlling of the incised valley. In the Gebel Ahmar Formation of the study area, several evidences point toward a tectonic inﬂuence on the ﬁlling of the incised valley. These include high sediment supply and deposition of braided river systems, transport of large clasts, and syn-sedimentary minor faults. The high sediment supply is manifested by the coarsegrained nature and vertical switching from sandy to gravelly braid rivers through the Oligocene in the study area (Fig. 14). The progradational pattern of the incised valley ﬁll succession, unusual for an incised valley, relate to the particular geological setting and to the interaction between tectonics and eustasy. By comparison of the sea level cycles of the nearby wells in southern Nile Delta, the erosional ﬂoor of the Oligocene incised valley, as a subaerial unconformity, is considered to be a sequence boundary ( Catuneanu, 2006). The lower part of the valley-ﬁll, above the sequence boundary, comprises ca 20 m of basal mainly sandy braided ﬂuvial deposits (storeys 1-2, Figs. 3, 14 A and B) accumulated during the base-level rise when sediment supply was sufﬁcient to outpace the rate of accommodation development on the shelf (lowstand to transgressive systems tract). Whereby the third phase of shoreline progradation resulting in a prominent progradation due to shifting from sandy to gravelly braid plain deposition (Fig. 14C and D). The gravelly braid plain of storeys 34 that were formed as the normal regressive highstand systems tract and eventually ﬁlled the valley. This progradation pattern reﬂects the increase in the sediments supply. The higher sediment

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supply is also conﬁrmed by the ﬂat to descending shelf edge trajectories in the Upper Oligocene (Selim, 2016). Although this study focuses on the exposures of the Oligocene potential reservoir, similar depositional analogs can be extrapolated to subsurface Oligocene reservoir rocks in the Cairo-Suez province. Four wells located on the north of Cairo-Suez province, for example, penetrated Oligocene incised valley ﬁlled with ﬂuvial strata, similar to those documented from this study (e.g. Selim, 2016). The architectural geometries of evolution of the braid plain system of this study can serve as direct analog for downdip and lateral equivalents for the Oligocene in the Cairo-Suez province and Nile delta. Slope channels formed downdip of both Oligocene lowstand and highstand ﬂuvial and alluvial deposits and are important in the emerging giant ﬁeld trends in the Oligocene in the onshore and offshore Nile Delta and Levant Basin (Dolson et al., 2005). The ﬂuvial incised valleys carry large volumes of eroded sediments toward the north not only in East Mediterranean basin but also into Arabian Gulf (Avni et al., 2012). The understanding of the Oligocene ﬂuvial incised valleys at North Eastern Desert helps to understand the architectures of the Oligocene deep water giant gas reservoirs in the onshore and offshore Nile Delta. It also represents a good documentation for shelf edge deltas that formed in similar tectonic regimes such as Matruh (NW Egypt, Tari et al., 2012), Levant and Cyprus Basins in the East Mediterranean ( Bowman, 2011; Gardosh and Druckman, 2006). In the Matruh basin, the ﬂuvial system sourced from the south feeds the Matruh canyon that ﬁlled with ponded clastics on shelf and slope (Tari et al., 2012). The tectonic inversion enhanced the sediment supply rate by Late Eocene-Oligocene. In the Levant Basin, where the Oligocene ﬂuvial valleys derived the sediments from the uplifted structures along Red Sea (Avni et al., 2012) and inverted structures (Gardosh and Druckman, 2006 ). These valleys linked to submarine canyons that ﬁlled with turbiditc deposits and distally grades into stacked channels and a basin ﬂoor fan (Gardosh and Druckman, 2006). The Cyprus Basin in offshore Syria contains up to 2900 m of Oligocene and Neogene sediments that are similar to that of Levant Basin (Bowman, 2011). The Oligocene sandstones in Palmyride region are common linked to tectonic uplift (Yzbek, 1998).

6. Conclusion The Oligocene Gebel Ahmar Formation represents up to 35 m deep incised valleys ﬁlled with aggrading ﬂuvial deposits. The valley conﬁguration is inferred from its geometry, multistorey internal architecture, size and ﬁlled with a ﬂuvial complex, underlined by a sequence boundary cutting into Middle to Upper Eocene marine carbonate bedrock. The valley basal surface is a sequence boundary, inasmuch as valley incision was promoted by a relative sea-level drop, whereas its ﬂuvial ﬁll was not only driven by a relative sea-level rise but also resulted from an increase in sediment supply due to the interplay between humid/monsoonal climate and hinterland tectonics. The valley ﬁlls consist of up to four vertically stacked storeys separated by ﬁfth-order boundingsurfaces. Each storey consists of multilateral primary channels and are interpreted as sandy and gravelly braid plain systems characterized by high, but ﬂuctuating, discharges and large sediment supplies. The dominant architectural elements in-channel include in braid plain gravel sheets and bars, gravel-sand couplets, gravel-sand foresets, downstream accretion DA, Downstreamoblique DLA, lateral accretion LA, channel (CH), sandy bedforms (SB), and Hollow HO. No ﬁne-grained overbank elements are preserved. The alluvial/ﬂuvial fan includes conglomerate-sandstone foresets, scour pool ﬁlling, and conglomerate-sand sheets architectural elements. The sediments were derived primarily from Eocene, Cretaceous, and Paleozoic strata, which were exposed uplifted northern Galala-Araba to the south. The initiation of the Suez rifting during the Late Oligocene caused widening and deepening of the ﬂuvial channels especially in storey 3. Acknowledgements The author is grateful to Editor Gregory Price, reviewer Martin stokes, and an anonymous reviewer for their critical comments and constructive criticism. The author thanks Prof. M. Darwish and Prof. A. El Manawi (Cairo University) for their support and critical discussion. Prof. Andrew Miall (University of Toronto) and Prof. Hugo Beraldi-Campesi (Arizona State University) are thanked for reviewing and editing the early version of this manuscript. References

5.3. Paleoclimate and vegetation implications As discussed above, the local climate is believed to inﬂuence the development of the incised-valley ﬁll. The abundance of crossbedding suggests a relatively constant discharge regime, indicating perennial ﬂuvial channels (e.g. Miall, 1996; Allen et al., 2013) that commonly found in humid environments (e.g. Almedeij and Diplas, 2005). This agrees with the conclusion of El-Kammar et al. (2000), who indicates the dominance of humid climate with heavy rains during the Oligocene. The presence of intrabasinal calcretes in storeys 2 and 3 (Fig.9E,F) indicates drier periods during calcrete formation (e.g. Gómez-Gras and Alonso-Zarza, 2003). The paleosols were sparsely covered by vegetation as revealed by the presence of large rhizoliths (Fig. 8E). Other evidence of the vegetation is represented by tree trunks described by Cuvillier (1930) and others. Rhizoliths, termites, wetland plants, coastal mangroves, and large trees of the Jebel Qatrani Formation indicate collectively that deposition took place under wet conditions, where tropical forest probably experienced a monsoonal climate (Bown, 1982; Wing and Tiffney, 1982; Olson and Rasmussen, 1986; Bown and Kraus, 1988; Rasmussen et al., 2001). In a global context, overall climatic change during the Oligocene with shifting from dominantly tropical to subtropical Eocene forests to more “mixed mesophytic forests” and ﬂoras adapted to (seasonal) dry conditions, followed by warm climates during the Middle to Late Oligocene (Miller et al., 1989).

Abd-Allah, A.M.A., 2008. Mesozoic-Cenozoic inversion tectonic of North Sinai: integration of structural and basin analysis. Journal of Applied Geophysics 7, 77– 108. Abdullatif, O.M., 1989. Channel-ﬁll and sheet-ﬂood facies sequences in the ephemeral-terminal River Gash, Kassala, Sudan. Sedimentary Geology 63, 171– 184. Aitken, J.F., Flint, S.S., 1994. High-frequency sequences and the nature of incised valley ﬁlls in ﬂuvial systems of the Breathitt Group (Pennsylvanian), Appalachian Foreland Basin, Eastern Kentucky. In: Dalrymple, R.W., Boyd, R., Zaitlin, B.A. (Eds.), Incised-Valley Systems: Origin and Sedimentary Sequences, 51. SEPM, Special Publication, pp. 352–368. Aldinucci, M., Ghinassi, M., Sandrelli, F., 2007. Climatic and tectonic signature in the ﬂuvial inﬁll of a late Pliocene Valley (Siena Basin, Northern Apennines, Italy). Journal of Sedimentary Research 77, 398–414. Alexander, J., Bridge, J.S., Leeder, M.R., Collier, R.E.L., Gawthorpe, R.L., 1994. Holocene meander-belt evolution in an active extensional basin: southwestern Montana. Journal of Sedimentary Research 64, 542–559. Allen, P.A., 1981. Sediments and processes on a small stream-ﬂow dominated, Devonian alluvial fan, Shetland Islands. Sedimentary Geology 29, 31–66. Allen, J.P., Fielding, C.R., Rygel, M.C., Gibling, M.R., 2013. Deconvolving signals of tectonic and climatic controls from continental basins: an example from the Late Paleozoic Cumberland Basin, Atlantic Canada. Journal of Sedimentary Research 83, 847–872. Allen, J.R.L., 1983. Studies in ﬂuviatile sedimentation: bars, bar-complexes and sandstone sheets (low-sinuosity braided streams) in the Brownstones (L. Devonian), Welsh Borders. Sedimentary Geology 33, 237–293. Allen, J.R.L., Williams, B.P.J., 1979. Interﬂuvial drainage on Siluro–Devonian alluvial plains in Wales and the Welsh borders. Journal of Geological Society 136, 361– 366. Almedeij, J., Diplas, P., 2005. Bed load sediment transport in ephemeral and perennial gravel bed streams. EOS, Transactions American Geophysical Union 86, 429–434.

S.S. Selim / Proceedings of the Geologists’ Association 128 (2017) 234–255 Anderson, M.P., Aiken, J.S., Webb, E.K., Mickelson, D.M., 1999. Sedimentology and hydrogeology of two braided stream deposits. Sedimentary Geology 129, 189– 199. Ardies, G.W., Dalrymple, R.W., Zaitlin, B.A., 2002. Controls on the geometry of incised valleys in the Basal Quartz Unit (Lower Cretaceous), western Canada sedimentary basin. Journal of Sedimentary Research 72, 602–618. Arnott, R.W.C., Zaitlin, B.A., Potocki, D.J., 2000. Geological controls on reservoir distribution in the Lower Cretaceous Basal Quartz, ChinCoulee–Horseﬂy Lake area, south-central Alberta. Bulletin of Canadian Petroleum Geology 48, 212– 229. Ashley, G.M., 1990. Classiﬁcation of large-scale subaqeous bedforms: a new look at an old problem. Journal of Sedimentary Petrology 60, 160–172. Ashmore, P.E., 1982. Laboratory modelling of gravel braided stream morphology. Earth Surface Processes and Landforms 7, 201–225. Ashmore, P.E., 1991. How do gravel-bar rivers braid? Canadian Journal of Earth Sciences 28, 326–341. Ashmore, P., 1993. Anabranch conﬂuences kinetics and sedimentation processes in gravel-braided streams. In: Best, J.B., Bristow, C.S. (Eds.), Braided Rivers, 75. Geological Society Publishing House, Bath, pp. 129–146 Special Publication. Ashworth, P.J., Best, J.L., Roden, J.E., Bristow, C.S., Klaassen, G.J., 2000. Morphological evolution and dynamics of a large, sand braid-bar, Jamuna River, Bangladesh. Sedimentology 47, 533–555. Ashworth, P.J., Sambrook Smith, G.H., Best, J.L., Bridge, J.S., Lane, S.N., Lunt, I.A., Reesink, A.J.H., Simpson, C.J., Thomas, R.E., 2011. Evolution and sedimentology of a channel ﬁll in the sandy braided South Saskatchewan River and its comparison to the deposits of an adjacent compound bar. Sedimentology 58, 1860–1883. Avni, Y., Segev, A., Ginat, H., 2012. Oligocene regional denudation of the northern Afar dome: Pre-and syn-breakup stages of the Afro-Arabian plate. Geological Society of America Bulletin 124, 1871–1897. Beadnell, H.J.L., 1905. The Topography and Geology of the Fayum Province of Egypt. Survey Department of Egypt, Cairo 101 p.. Best, J.L., 1988. Sediment transport and bed morphology at river channel conﬂuences. Sediementology 35, 481–498. Best, J.L., Ashworth, P.J., 1997. Scour in large braided rivers and the recognition of sequence stratigraphic boundaries. Nature 387, 275–277. Best, J.L., Ashworth, P.J., Bristow, C.S., Roden, J., 2003. Three-dimensional sedimentary architecture of a large, mid-channel sand braid bar, Jamuna River, Bangladesh. Journal of Sedimentary Research 73, 516–530. Bianchi, V., Salles, T., Ghinassi, M., Billi, P., Dallanave, E., Duclaux, G., 2015. Numerical modeling of tectonically driven river dynamics and deposition in an upland incised valley. Geomorphology 241, 353–370. Blair, T.C., McPherson, J.G., 1994. Alluvial fans and their natural distinction from rivers based on morphology, hydraulic processes, sedimentary processes, and facies assemblages. Journal of Sedimentary Research 64, 450–489. Blum, M.D., 1992. Climatic and eustatic controls on Gulf coastal plain sedimentation: an example from the late Quaternary of the Colorado River. In: Armentrout, J.M., Perkins, R.F. (Eds.), Sequence Stratigraphy as an Exploration Tool: Concepts and Practices in the Gulf Coast: SEPM, Gulf Coast Section, Eleventh Annual Research Conference, Houston, pp. 71–84. Blum, M.D., Aslan, A., 2006. Signatures of climate vs. sea-level change within incised valley-ﬁll successions: quaternary examples from the Texas Coastal Plain. Sedimentary Geology 190, 177–211. Blum, M.D., Price, D.M., 1998. Quaternary alluvial plain construction in response to glacio-eustatic and climatic controls, Texas Gulf coastal plain. In: Shanley, K.W., McCabe, P.J. (Eds.), Relative Role of Eustasy, Climate, and Tectonism in Continental Rocks, 59. SEPM, Special Publication, pp. 31–48. Boothroyd, J.C., Nummedal, D., 1978. Proglacial braided outwash: a model for humid alluvial-fan deposits. Canadian Society of Petroleum Geologists Memoir 5, 641– 668. Bowman, S.A., 2011. Regional seismic interpretation of the hydrocarbon prospectivity of offshore Syria. GeoArabia 16, 95–124. Breda, A., Amorosi, A., Rossi, V., Fusco, F., 2016. Late-glacial to Holocene depositional architecture of the Ombrone palaeovalley system (Southern Tuscany, Italy): sealevel, climate and local factors control in valley-ﬁll variability. Sedimentology 63, 1124–1148. doi:http://dx.doi.org/10.1111/sed.12253. Bristow, C.S., 1987. Brahmaputra River: channel migration and deposition. In: Ethridge, F.G., Flores, R.M., Harvey, M.D. (Eds.), Recent Developments in Fluvial Sedimentology, 39. Society of Economic Paleontologists and Mineralogists Special Publication, Tulsa, OK, pp. 63–74. Bridge, J.S., Gabel, S.L., 1992. Flow and sediment dynamics in a low sinuosity, braided river: Calamus River, Nebraska Sandhills. Sedimentology 39, 125–142. Bridge, J.S., Alexander, J., Collier, R.E.L., Gawthorpe, R.L., Jarvis, J., 1995. Groundpenetrating radar and coring used to study the large-scale structure of point-bar deposits in three dimensions. Sedimentology 42, 839–852. Bridge, J., Collier, R., Alexander, J., 1998. Large-scale structure of Calamus River deposits (Nebraska, USA) revealed using ground-penetrating radar. Sedimentology 45, 977–986. Boothroyd, J.C., Ashley, G.M., 1975. Processes, bar morphology, and sedimentary structures on braided outwash fans, northeastern Gulf of Alaska. In: Jopling, A. V., McDonald, B.C. (Eds.), Glacioﬂuvialand Glaciolacustrine Sedimentation, 23. SEPM Spec. Publ., pp. 193–222. Bosworth, W., McClay, K., 2001. Structural and stratigraphic evolution of the Gulf of Suez Rift, Egypt, a synthesis (in Peri-Tethys memoir 6, Peri-Tethyan Rift/Wrench basins and passive margins). Memoires du Museum National d’Histoire Naturelle 186, 567–606.

253

Bown, T.M., 1982. Ichnofossils and rhizoliths of the nearshore ﬂuvial Jebel Qatrani Formation (Oligocene), Fayum Province, Egypt. Palaeogeography, Palaeoclimatology, Palaeoecology 40, 255–309. Bown, T.M., Kraus, M.J., 1988. Geology and paleoenvironment of the Oligocene Jebel Qatrani Formation and adjacent rocks, Fayum Depression, Egypt, 1452. United States Geological Survey Professional Paper 64 p. Boyd, R., Dalrymple, R.W., Zaitlin, B.A., 2006. Estuarine and incised-valley facies models, 84. SEPM Special Publication, pp. 171. Bridge, J.S., 2003. Rivers and Floodplains. Blackwell Scientiﬁc, Oxford 504 pp.. Brierley, G.J., Liu, K., Crook, K.A.W., 1993. Sedimentology of coarse-grained alluvial fans in the Markham Valley, Papua New Guinea. Sedimentary Geology 86, 297– 324. Bristow, C.S., 1993. Sedimentary structures exposed in bar tops in the Brahmaputra River, Bangladesh. In: Best, J.L., Bristow, C.S. (Eds.), Braided Rivers, 75. Geological Society of London Special Publication 419 pp.. Buatois, L.A., Mángano, M.G., 2004. Applications of ichnology in facies and sequence stratigraphic Animal-substrate interactions freshwater environments: analysis of ﬂuvio-lacustrine successions. In: McIlroy, D. (Ed.), The Application of lchnology to Palaeoenvironmental and Stratigraphic Analysis, 228. Geological Society, London, Special Publications, pp. 311–333. olaza, F.G., 2002. Trazas fósiles: Sen ales de Buatois, L.A., Mángano, M.G., Acen comportamiento en el registro estratigráﬁco, 2. Museo Paleontológico Egidio Feruglio Edición Especial. Campbell, C.V., 1976. Reservoir geometry of a ﬂuvial sheet sandstone. AAPG Bulletin 60, 1009–1020. Cant, D.J., Walker, R.G., 1978. Fluvial processes and facies sequences in the sandy braided South Saskatchewan River, Canada. Sedimentology 25, 625–648. Catuneanu, O., 2006. Principles of Sequence Stratigraphy. Elsevier, Amsterdam 386 pp.. Chamyal, L.S., Khadkikar, A.S., Malik, J.N., Maurya, D.M., 1997. Sedimentology of the Narmada alluvial fan, western India. Sedimentary Geology 107, 263–279. Clarke, M.R., Dixon, A.J., 1981. The Pleistocene braided river deposits in the Blackwater Valley area of Berkshire and Hampshire, England. Proceedings of the Geologists’ Association 92, 139–157. Cole, J.M., 1996. Aggradation, degradation, and levels of heterogeneity of the braided Niobrara River, NE Nebraska. Unpublished MS thesis. Colorado State University, Fort Collins 131 pp. Coleman, J.M., 1969. Brahmaputra River: channel processes and sedimentation. Sedimentary Geology 3, 129–239. Colombera, L., Mountney, N.P., McCaffrey, W.D., 2012a. A relational database for the digitization of ﬂuvial architecture: concepts and example applications. Petroleum Geoscience 18, 129–140. Colombera, L., Felletti, F., Mountney, N.P., McCaffrey, W.D., 2012b. A database approach for constraining stochastic simulations of the sedimentary heterogeneity of ﬂuvial reservoirs. The American Association of Petroleum Geologists Bulletin 96, 2143–2166. Colombera, L., Mountney, N.P., McCaffrey, W.D., 2013. A quantitative approach to ﬂuvial facies models: methods and example results. Sedimentology 60, 1526– 1558. Collinson, J.D., 1978. Vertical sequences and sand body shape in alluvial sequences. Canadian Society of Petroleum Geologists Memoir 5, 577–586. Collinson, J.D., 1996. Alluvial sediments, In: Reading, H.G. (Ed.), Sedimentary Environments and Facies. 3rd edn. Blackwell Science, Oxford, pp. 37–82. Collinson, J.D., Mountney, N.P., Thompson, D.B., 2006. Sedimentary Structures, 3rd edn Terra Publishing, Harpenden 292 pp.. Cuvillier, J., 1930. Revision du Nummulitique Egyptien. Mémoires de l'Institut d'Egypte 16, 371. Dalrymple, R.W., Boyd, R., Zaitlin, B.A., 1994. History of research, types and internal organization of incised-valley systems: introduction to the volume. In: Dalrymple, R.W., Boyd, R., Zaitlin, B.A. (Eds.), Incised-Valley Systems: Origin and Sedimentary Sequences, 51. SEPM, Special Publication, pp. 3–10. Dalrymple, M., 2010. Interpreting sedimentary successions: facies, facies analysis and facies models. In: James, N.P., Dalrymple, R.W. (Eds.), Facies Models 4, 6. Geological Association of Canada, St. John's, Newfoundland, pp. 3–18 GEO text. Dolson, J.C., El Barkooky, A.N., Wehr, F., Gingerich, P., Prochazka, N., Shann, M., 2002. The Eocene and Oligocene Paleo-Ecology and Paleo-Geography of Whale Valley and the Fayoum Basins: Implications for Hydrocarbon Exploration in the Nile Delta and Eco-Tourism in the Greater Fayoum Basin. AAPG, Search and Discovery Article #10030. http://www.searchanddiscovery.com/documents/ cairo/index.htm. Dolson, J.C., Boucher, P.J., Siok, J., Heppard, P.D., 2005. Key challenges to realizing full potential in an emerging giant gas province: Nile Delta/Mediterranean offshore, deep water, Egypt. Geological Society, London, Petroleum Geology Conference series 6, 607–624. Dolson, J.C., 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, 106. American Association of Petroleum Geologists Memoir, pp. 143–178. El-Kammar, A., El-Kammar, M., El-Gamal, I., 2000. Rare earth elements and uranium geochemistry of terrestrial Tertiary bones from Fayium, Egypt. 5th International Conference on the Geology of the Arab World 213–238. El Naggar, S., El Morshedy, A., 2013. Regional study of Oligocene Future Hydrocarbon Potential in Nile Delta, Egypt. SEG Annual Meeting, 22–27 September, Houston, Texas, conference paper. Evans, J.E., 1991. Facies relationships, alluvial architecture, and paleohydrology of a Paleogene, humid-tropical alluvial-fan system: Chumstick Formation, Washington State, USA. Journal of Sedimentary Petrology 61, 732–755.

254

S.S. Selim / Proceedings of the Geologists’ Association 128 (2017) 234–255

Fitzgerald, P.G., Barrett, P.J., 1986. Skolithos in a Permian braided river deposit, southern Victoria Land, Antarctica. Palaeogeography, Palaeoclimatology, Palaeoecology 52, 237–247. Gabel, S.L., 1993. Geometry and kinematics of dunes during steady and unsteady ﬂows in the Calamus River, Nebraska, USA. Sedimentology 40, 237–269. Gardosh, M.A., Druckman, Y., 2006. Seismic Stratigraphy, Structure and Tectonic Evolution of the Levantine Basin, 260. Geological Society, London, pp. 201–227 Special Publications. Geehan, G., Underwood, J., 1993. The use of length distributions in geological modelling. In: Flint, S.S., Bryant, I.D. (Eds.), The Geological Modelling of Hydrocarbon Reservoirs and Outcrop Analogues, 15. Int. Assoc. Sedimentol. Spec. Publ., pp. 205–212. Gibling, M.R., 2006. Width and thickness of ﬂuvial channel bodies and valley ﬁlls in the geological record: a literature compilation and classiﬁcation. Journal of Sedimentary Research 76, 731–770. Gómez-Gras, D.Y., Alonso-Zarza, A.M., 2003. Reworked calcretes: their signiﬁcance in the reconstruction of alluvial sequences (Permian and Triassic, Minorca, Balearic Islands, Spain). Sedimentary Geology 158, 299–319. Guiraud, R., Bosworth, W., Thierry, J., Delplanque, A., 2005. Phanerozoic geological evolution of Northern and Central Africa: an overview. Journal of African Earth Sciences 43, 83–143. Haq, B.U., Al-Qahtani, A.M., 2005. Phanerozoic cycles of sea-level change on the Arabian Platform. GeoArabia 10 (2), 127–159. Haq, B.U., Hardenbol, J., Vail, P.R., 1988. Mesozoic and Cenozoic chronostratigraphy and cycles of sealevel change. Society of Economic Paleontologists and Mineralogists 42, 71–108. Hein, F.J., Walker, R.G., 1977. Bar evolution and development of stratiﬁcation in the gravelly, braided, Kicking Horse River, British Columbia. Canadian Journal of Earth Science 14, 562–570. Hofmann, C., Courtillot, V., Feraud, G., Rochette, P., Yirgu, G., Ketefo, E., Pik, R., 1997. Timing of the Ethiopian ﬂood basalt event and implications for plume birth and global change. Nature 389, 838–841. Holbrook, J., 2001. Origin, genetic interrelationships, and stratigraphy over the continuum of ﬂuvial channel-form bounding surfaces: an illustration from middle Cretaceous strata, southeastern Colorado. Sedimentary Geology 144, 179–222. Huggenberger, P., 1993. Radar facies: recognition of characteristic braided river structures of the Pleistocene Rhine gravel (NE part of Switzerland). In: Best, J.B., Bristow, C.S. (Eds.), Braided Rivers, 75. Geological Society Publishing House, Bath, pp. 163–176 Special Publication. Hussein, I.M., Abd-Allah, A.M.A., 2001. Tectonic evolution of the northeastern part of the African continental margin, Egypt. Journal of African Earth Sciences 33, 49– 68. Issawi, B., McCauley, J.F., 1993. The Cenozoic landscape of Egypt and its river systems. Annals of the Geological Survey of Egypt 19, 357–384. Jensen, M.A., Pedersen, G.K., 2010. Architecture of vertically stacked ﬂuvial deposits, Atane Formation, Cretaceous, Nuussuaq, central West Greenland. Sedimentology 57, 1280–1314. Jo, H.R., Rhee, C.W., Chough, S.K., 1997. Distinctive characteristics of a streamﬂowdominated alluvial fan deposit: Sanghori area, Kyongsang Basin (Early Cretaceous), southeastern Korea. Sedimentary Geology 110, 51–79. Jones, S.J., Frostick, L.E., Astin, T.R., 2001. Braided stream and ﬂood plain architecture: the Rio Vero Formation, Spanish Pyrenees. Sedimentary Geology 139, 229–260. Ismail, A.A., Abdel Ghany, O., 1999. Lower Miocene foraminifera from some exposures in the Cairo-Suez district, Eastern Desert, Egypt. Journal of African Earth Sciences 26 (3), 507–526. Kleinspehn, K.L., Steel, R.J., Johannessen, E., Netland, A., 1984. Conglomeratic fandelta sequences, Late CarboniferousEarly Permian, western Spitsbergen. In: Koster, E.H., Steel, R.J. (Eds.), Sedimentology of Gravels and Conglomerates, 10. Canadian Society of Petroleum Geologists memoir, pp. 279–294. Klitzsch, E., List, F. K., Pohlmann, G. Handley, R. Hermina, M. and Meissner, G. eds., 1986, 1987, Geological map of Egypt: 1:50,000 scale, 20 sheets, Cairo, Conoco/ Egyptian General Petroleum Corporation. Kostic, B., Aigner, T., 2007. Sedimentary architecture and 3D ground-penetrating radar analysis of gravelly meandering river deposits (Neckar Valley, SW Germany). Sedimentology 54, 789–808. Kuss, J., Scheibner, C., Gietl, R., 2000. Carbonate platform to basin transition along an upper Cretaceous to lower tertiary Syrian Arc uplift, Galala Plateaus, Eastern Desert, Egypt. GeoArabia 5, 405–424. Lahiri, S.K., Sinha, R., 2012. Tectonic controls on the morphodynamics of the Brahmaputra River system in the upper Assam valley, India. Geomorphology 169, 74–85. Leclair, S.F., Bridge, J.S., 2001. Quantitative interpretation of sedimentary structures formed by river dunes. Journal of Sedimentary Research 71, 713–716. Liu, K., Boult, P., Painter, S., Paterson, L., 1996. Outcrop analog for sandy braided stream reservoirs: permeability patterns in the Triassic Hawkesbury Sandstone, Sydney Basin, Australia. American Association of Petroleum Geologist Bulletin 80, 1850–1866. Long, D.G., 2006. Architecture of pre-vegetation sandy-braided perennial and ephemeral river deposits in the Paleoproterozoic Athabasca Group, northern Saskatchewan, Canada as indicators of Precambrian ﬂuvial style. Sedimentary Geology 190, 71–95. Lunt, I.A., Bridge, J.S., 2004. Evolution and deposits of a gravelly braid bar, Sagavanirktok River, Alaska. Sedimentology 51, 415–432.

Lunt, I.A., Bridge, J.S., Tye, R.S., 2004. A quantitative three-dimensional depositional model of gravelly braided rivers. Sedimentology 51, 377–414. McCabe, P.J., 1977. Deep distributary channels and giant bedforms in the Upper Carboniferous of the Central Pennines, northern England. Sedimentary Geology 24, 271–290. McLelland, S.J., Ashworth, R.J., Best, J.L., Roden, J., Klaassen, G.J., 1999. Flow structure and transport of sand-grade suspended sediment around an evolving braid bar, Jamuna River, Bangladesh. In: Smith, N.D., Rogers, J. (Eds.), Fluvial Sedimentology VI. Special Publication 28 International Association of Sedimentologists. Blackwell Science Oxford, pp. 43–57. Main, C.E., Robertson, A.H., Palamakumbura, R.N., 2016. Pleistocene geomorphological and sedimentary development of the Akaki River catchment (northeastern Troodos Massif) in relation to tectonic uplift versus climatic change. International Journal of Earth Sciences 105, 463–485. Maizels, J., 1993. Lithofacies variations within sandur deposits: the role of runoff regime, ﬂow dynamics and sediment supply characteristics. Sedimentary Geology 85, 299–325. Marren, P.M., 2001. Sedimentology of proglacial rivers in eastern Scotland during the Late Devensian. Earth and Environmental Science Transactions of the Royal Society of Edinburgh 92, 149–171. Marriott, S.B., Wright, V.P., 1993. Palaeosols as indicators of geomorphic stability in two Old Red Sandstone alluvial suites, South Wales. Journal Geological Society (London) 150, 1109–1120. Martin, J., Cantelli, A., Paola, C., Blum, M., Wolinsky, M., 2011. Quantitative modeling of the evolution and geometry of incised valleys. Journal of Sedimentary Research 81, 64–79. Mattheus, C.R., Rodriguez, A.B., 2011. Controls on late Quaternary incised-valley dimension along passive margins evaluated using empirical data. Sedimentology 1113–1137. McGowen, J.H., Groat, C.G., 1971. Van Horn sandstone, West Texas: an alluvial fan model mineral exploration. Bureau of Economic Geology. Report of Investigations 72 57 pp. Mess, A., Mohr, B., Martin, T., 2001. Evolutionary transformations of hystricognath Rodentia and the climatic change in the Eocene to Late Oligocene time interval. Zoosystematics and Evolution 77 (2), 193–206. Meulenkamp, J.E., Sissingh, W., 2003. Tertiary palaeogeography and tectonostratigraphic evolution of the Northern and Southern Peri-Tethys platforms and the intermediate domains of the African–Eurasian convergent plate boundary zone. Palaeogeography, Palaeoclimatology, Palaeoecology 196, 209–228. Miall, A.D., 1977. A review of the braided-river depositional environment. Earth Science Reviews 13, 1–62. Miall, A.D., 1978. Fluvial Sedimentology, 5. Canadian Society of Petroleum Geologists Memoir 859 p.. Miall, A.D., 1985. Architectural-element analysis: a new method of facies analysis applied to ﬂuvial deposits. Earth science Review 22, 261–308. Miall, A.D., 1988. Architectural elements and bounding surfaces in ﬂuvial deposits: anatomy of the Kayenta Formation (Lower Jurassic), southwest Colorado. Sedimentary Geology 55, 233–262. Miall, A.D., 1993. The architecture of ﬂuvial-deltaic sequences in the upper Mesaverde Group (Upper Cretaceous), Book Cliffs, Utah. Geological Society London, Special Publications 75, 305–332. Miall, A.D., 1994. Reconstructing ﬂuvial macroform architecture from twodimensional outcrops: examples from the Castlegate Sandstone, Book Cliffs, Utah. Journal of Sedimentary Research 64, 146–158. Miall, A.D., 1996. The Geology of Fluvial Deposits. Springer-Verlag, New York 582 p p. Miall, A.D., 2010a. Alluvial deposits. In: James, N.P., Dalrymple, R.W. (Eds.), Facies Models 4. Geol. Assoc., Canada, St. John’s, Newfoundland, pp. 105–137. Miall, A.D., 2010b. The Geology of Fluvial Deposits: Sedimentary Facies Basin Analysis and Petroleum Geology, 4th edn Springer, Berlin 582 pp. Miall, A.D., Arush, M., 2001. Cryptic sequence boundaries in braided ﬂuvial successions. Sedimentology 48, 971–985. Miall, A.D., 2014. Fluvial Depositional Systems. Springer, Switzerland 316 pp.. Miall, A.D., Jones, B.G., 2003. Fluvial architecture of the Hawkesbury sandstone (Triassic), near Sydney, Australia. Journal of Sedimentary Research 73, 531–545. Middleton, L.T., Trujillo, A.P., 1984. Sedimentology and depositional setting of the Upper Proterozoic Scanlan conglomerate, Central Arizona. In: Koster, E.H., Steel, R.J. (Eds.), Sedimentology of Gravels and Conglomerates, 10. Mem. Can. Soc. Petrol. Geol., Calgary, pp. 181–202. Miller, K.G., Wright, J.D., Brower, A.N., 1989. Oligocene to Miocene stable isotope stratigraphy and planktonic foraminifer biostratigraphy of the Sierra Leone Rise (DSDP Site 366 and ODP Site 667). In: Ruddiman, W., Sarnthein, M., et al. (Eds.), Proceedings of the Ocean Drilling Program, vol. 108. Scientiﬁc Results Ocean Drilling Program College Station, Texas, pp. 279–294. Miller, K.G., Browning, J.V., Aubry, M.P., Wade, B.S., Katz, M.E., Kulpecz, A.A., Wright, J.D., 2008. Eocene–Oligocene global climate and sea-level changes: St. Stephens Quarry, Alabama. Geological Society of America Bulletin 120, 34–53. Moustafa, A.R., Khalil, M.H., 1995. Superposed deformation in the northern Suez rift, Egypt: relevance to hydrocarbons exploration. Journal of Petroleum Geology 18, 245–266. Mosselman, E., Koomen, M.H.E., Seijmonsbergen, A.C., 1995. Morphological changes in a large braided sand-bed river. In: Hickin, E.J. (Ed.), River Geomorphology. Wiley, Chichester, pp. 235–249. Mumpy, A.J., Jol, H.M., Kean, W.F., Isbell, J.L., 2007. Architecture and sedimentology of an active braid bar in the Wisconsin River based on 3-D ground penetrating

S.S. Selim / Proceedings of the Geologists’ Association 128 (2017) 234–255 radar. In: Barker, G.S., Jol, H.M. (Eds.), Stratigraphic Analysis Using GPR, 432. Geological Society of America, Special Paper, pp. 111–131. Nemec, W., Postma, G., 1993. Quaternary alluvial fans in southwestern Crete: sedimentation processes and geomorphic evolution. In: Marzo, M., Puigdefáibregas, C. (Eds.), Alluvial sedimentation, 17. International Assocation of Sedimentologists Special Publication, pp. 235–276. Nichols, G.J., Hirst, J.P., 1998. Alluvial fans and ﬂuvial distributary systems OligoMiocene, Northern Spain: contrasting processes and products. Journal of Sedimentary Research 68, 879–889. Olson, S.L., Rasmussen, D.T., 1986. Paleoenvironment of the earliest hominoids; new evidence from the Oligocene avifauna of Egypt. Science 233, 1202–1204. Ori, G.G., 1982. Braided to meandering channel patterns in humid-region alluvial fan deposits, River Reno, Po Plain (northern Italy). Sedimentary Geology 31, 231– 248. Owen, G., 1996. Experimental soft-sediment deformation: structures formed by the liquefaction of unconsolidated sands and some ancient examples. Sedimentology 43, 279–293. Patton, T.L., Moustafa, A.R., Nelson, R.A., Abdine, A.S., 1994. Tectonic evolution and structural setting of the Suez Rift (in Interior Rift Basins, Landon). AAPG Memoirs 59, 9–55. Ramos, A., Sopeña, A., 1983. Gravel bars in low-sinuosity streams (Permian and Triassic, central Spain). In: Collinson, J.D., Lewin, J. (Eds.), Modern and Ancient Fluvial Systems, 6. International Association of Sedimentologists, Special Publication, pp. 301–312. Rasmussen, E.S., 2014. Development of an incised-valley ﬁll under the inﬂuence of tectonism and Glacio-Eustatic sea-level change: valley Morphology, Fluvial Style, and Lithology. Journal of Sedimentary Research 84, 278–300. Rasmussen, D.T., Simons, E.L., Hertel, F., Judd, A., 2001. Hindlimb of a giant terrestrial bird from the Upper Eocene, Fayum, Egypt. Palaeontology 44 (2), 325–337. Reid, I., Frostick, L.E., 1987. Towards a better understanding of bedload transport. In: Ethridge, F.G., Flores, R.M., Harvey, M.D. (Eds.), Recent Developments in Fluvial Sedimentology, 39. Soc. Econ. Paleontol. Mineral. Special Publication, pp. 13–20. Reesink, A.J., Bridge, J.S., 2009. Inﬂuence of bedform superimposition and ﬂow unsteadiness on the formation of cross strata in dunes and unit bars-Part 2, further experiments. Sedimentary Geology 222, 274–300. Reesink, A.J.H., Ashworth, P.J., Sambrook Smith, G.H., Best, J.L., Parsons, D.R., Amsler, M.L., Hardy, R.J., Lane, S.N., Nicholas, A.P., Orfeo, O., Sandbach, S.D., Simpson, C.J., Szupiany, R.N., 2014. Scales and causes of heterogeneity in bars in a large multichannel river: Río Paraná, Argentina. Sedimentology 61, 1055–1085. Richardson, W.R., Thorne, C.R., 2001. Multiple thread ﬂow and channel bifurcation in a braided river: Brahmaputra–Jamuna River, Bangladesh. Geomorphology 38, 185–196. Richardson, W.R.R., Thorne, C.R., Mahmood, S., 1996. Secondary ﬂow and channel changes around a bar in the Brahmaputra River, Bangladesh. In: Ashworth, P.J., Bennett, S.J., Best, J.L., McLelland, S.J. (Eds.), Coherent Flow Structures in Open Channels. Wiley, Chichester, pp. 520–543. Rodriguez, A.B., Anderson, J.B., Bradford, J., 1998. Holocene tidal deltas of the Trinity incised valley: analogs for exploration and production. Gulf Coast Association of Geological Societies Transactions 48, 373–380. Rögl, F., 1998. Palaeogeographic considerations for Mediterranean and Paratethys Seaways (Oligocene to Miocene). Annalen des Naturhistorischen Museums in Wien 99A, 279–310. Røe, S.L., Hermansen, M., 2006. New aspects of deformed cross-strata in ﬂuvial sandstones: examples from Neoproterozoic formations in northern Norway. Sedimentary Geology 186, 283–293. Rust, B.R., 1972. Structure and process in a braided river. Sedimentology 18, 221–245. Rust, B.R., 1978. Depositional models for braided alluvium. In: Miall, A.D. (Ed.), Fluvial Sedimentology, 5. Canadian Society of Petroleum Geology Memoir, pp. 605–625. Rust, B.R., 1979. Coarse alluvial deposits. In: Walker, R.G. (Ed.), Facies Models, 1. Geoscience Canada Reprint Series, pp. 9–21. Sadek, H., 1926. The geography and geology of the district between Gebel Ataqa and El Galala El Bahariya (Gulf of Suez). Geological Survey of Egypt 1–120. Said, R., 1981. The Geological Evolution of the River Nile. Springer, Heidelberg 151 p.. Sambrook Smith, G.H., Ashworth, P.J., Best, J.L., Woodward, J., Simpson, C.J., 2005. The morphology and facies of sandy braided rivers: some considerations of spatial and temporal scale invariance. In: Blum, M.D., Marriott, S.B., Leclair, S.F. (Eds.), Fluvial Sedimentology VII, 35. International Association of Sedimentologists, Special Publication, pp. 145–158. Sambrook Smith, G.H., Ashworth, P.J., Best, J.L., Woodward, J., Simpson, C.J., 2006. The sedimentology and alluvial architecture of the sandy braided South Saskatchewan River, Canada. Sedimentology 53, 413–434. Sarker, M.H., Thorne, C.R., Aktar, M.N., Ferdous, M.R., 2014. Morpho-dynamics of the Brahmaputra–Jamuna River, Bangladesh. Geomorphology 215, 45–59. Schumm, S.A., 1993. River response to baselevel change: implications for sequence stratigraphy. Journal of Geology 101, 279–294. Schumm, S.A., Mosley, M.P., Weaver, W.E., 1987. Experimental Fluvial Geomorphology. Wiley, New York.

255

Scherer, C.M., Goldberg, K., Bardola, T., 2015. Facies architecture and sequence stratigraphy of an early post-rift ﬂuvial succession, Aptian Barbalha Formation, Araripe Basin, northeastern Brazil. Sedimentary Geology 322, 43–62. Selim, S.S., 2016. Sedimentological architecture, shelf-edge trajectories and evolution of an Oligocene reservoir, East Nile Delta. Geological Magazine accepted manuscript. Shukri, N.M., 1954. Remarks on the geological structure of Egypt. Geographical Society of Egypt Bulletin 2, 1–19. Siegenthaler, C., Huggenberger, P., 1993. Pleistocene Rhine gravel: deposits of a braided river system with dominant pool preservation. In: Best, J.B., Bristow, C.S. (Eds.), Braided Rivers, 75. Geological Society Publishing House, Bath, pp. 147– 162 Special Publication. Singh, A., Bhardwaj, B.D., 1991. Fluvial facies model of the Ganga River sediments, India. Sedimentary Geology 72, 135–146. Skelly, R.L., Bristow, C.S., Ethridge, F.G., 2003. Architecture of channel-belt deposits in an aggrading shallow sandbed braided river: the lower Niobrara River northeast Nebraska. Sedimentary Geology 158, 249–270. Smith, N.D., 1970. The braided stream depositional environment: comparison of the Platte River with some Silurian Clastic Rocks, North Central Appalachians. Geological Society of America Bulletin 81, 2993–3014. Smith, N.D., 1971. Transverse bars and braiding in the Lower Platte River, Nebraska. Geological Society of America Bulletin 82, 3407–3420. Smith, N.D., 1972. Some sedimentological aspects of planar cross-stratiﬁcation in a Sandy Braided River. Journal of Sedimentary Petrology 42, 624–634. Smith, G.H.S., Ashworth, P.J., Best, J.L., Lunt, I.A., Orfeo, O., Parsons, D.R., 2009. The sedimentology and alluvial architecture of a large braid bar, Río Paraná, Argentina. Journal of Sedimentary Research 79, 629–642. Smith, N.D., 1974. Sedimentology and bar formation in the upper Kicking Horse River, a braided outwash stream. Journal of Geology 82, 205–224. Smith, S.A., 1990. The sedimentology and accretionary styles of an ancient gravelbed stream; the Budleigh Salterton pebble beds (Lower Triassic), Southwest England. Sedimentary Geology 67 (3–4), 199–219. Srivastava, P., Ray, Y., Phartiyal, B., Sharma, A., 2013. Late Pleistocene-Holocene morphosedimentary architecture, Spiti River, arid higher Himalaya. International Journal of Earth Sciences 102 (7), 1967–1984. Soltan, R., Mountney, N.P., 2016. Interpreting complex ﬂuvial channel and barform architecture: Carboniferous Central Pennine Province, northern England. Sedimentology 63, 207–252. Steel, R.J., Thompson, D.B., 1983. Structures and textures in Triassic braided stream conglomerates (“Bunter” pebble beds) in the Sherwood Sandstone Group, North Staffordshire, England. Sedimentology 30, 341–367. Strougo, A., Abd-Allah, A.M., 1990. Mokattamian stratigraphy of northcentral Eastern Desert (south of Maadi-Qattamiya road) Middle East Research Center, Ain Shams University. Earth Science Series 4, 152–175. Strougo, A., Bignot, G., Abd-Allah, A.M., 1992. Biostratigraphy and paleoenvironments of Middle Eocene Benthic foraminiferal assemblages of northcentral Eastern Desert Egypt, Middle East Research Center, Ain Shams University. Earth Science Series 6, 1–12. Sundborg, Å., 1956. The River Klaralven: a study of ﬂuvial processes. Geograﬁska Annaler 38 (3), 238–316. Tari, G., Hussein, H., Novotny, B., Hannke, K., Kohazy, R., 2012. Play types of the deepwater Matruh and Herodotus basins, NW Egypt. Petroleum Geoscience 18, 443– 455. Thorne, C.R., Russell, A.P.G., Alam, M.K., 1993. Planform pattern and channel evolution of the Brahmaputra River, Bangladesh. In: Best, J.L., Bristow, C.S. (Eds.), Braided Rivers, 75. Geological Society of London, Special Publication, pp. 257–276. Todd, S.P., 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, 513–530. Todd, S.P., 1996. Process deduction from ﬂuvial sedimentary structures. In: Carling, P.A., Dawson, M.R. (Eds.), Advances in Fluvial Dynamics and Stratigraphy. Wiley, Chichester, pp. 299–350. Wang, Z., Yu, G., Wang, X., Melching, C.S., Liu, L., 2015. Sediment storage and morphology of the Yalu Tsangpo valley due to uneven uplift of the Himalaya. Science China Earth Sciences 6 p. Article in press. Willis, B.J., 1997. Architecture of ﬂuvial-dominated valley-ﬁll deposits in the Cretaceous Fall River Formation. Sedimentology 44, 735–757. Wing, S.L., Tiffney, B.H., 1982. A paleotropical ﬂora from the Oligocene Jebel Qatrani Formation of northern Egypt: a preliminary report. Miscellaneous Studies of the Botanical Society of America 162, 67. Wright, V.P., Marriot, S.B., 1993. The sequence stratigraphy of ﬂuvial depositional systems: the role of ﬂoodplain sediment storage. Sedimentary Geology 86, 203– 210. Wooldridge, C.L., Hickin, E.J., 2005. Radar architecture and evolution of channel bars in wandering gravel-bed rivers: Fraser and Squamish rivers, British Columbia, Canada. Journal of Sedimentary Research 75, 844–860. Yzbek, M.K., 1998. Calcareous nannofosil zonation and micofacies studies on the Eocene sediments of the Palmyridian and Anti-Lebanon region, Syria. Unpublished PhD thesis. Cairo University 342 p..

Facies and architecture of a coarse-grained alluvial-dominated incised valley fill: a case study from the Oligocene Gebel Ahmar Formation, Southern Tethyan-shelf (northern Egypt)

Facies and architecture of a coarse-grained alluvial-dominated incised valley fill: a case study from the Oligocene Gebel Ahmar Formation, Southern Tethyan-shelf (northern Egypt)

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