Disconnected submarine lobes as a record of stepped slope evolution over multiple sea-level cycles Open Access

Topographic complexity across submarine slopes varies both spatially and temporally (Fig. 1). Topography can include three-dimensional confinement of ponded mini-basins (e.g., Prather et al., 1998, 2012a, 2012b, 2017; Badalini et al., 2000; Winker and Booth, 2000; Sinclair and Tomasso, 2002; Shultz and Hubbard, 2005; Sylvester et al., 2015); barriers creating tortuous corridors (e.g., Smith, 2004a; Hay, 2012); and more subtle gradient changes (generally <1° to few degrees) of stepped slope profiles (Figs. 1B and 1C), which create higher gradient ramps linking lower gradient steps (e.g., O’Byrne et al., 2004; Smith, 2004a; Barton, 2012; Deptuck et al., 2012; Hay, 2012). High-amplitude seabed topography and associated major gradient changes have been well documented in seismic reflection data sets of modern systems, such as in the Gulf of Mexico (Prather et al., 1998, 2017; Fig. 1D), the Western Niger Delta slope (Jobe et al., 2017; Fig 1E), offshore Norway (Jackson et al., 2008) and offshore Angola (Hay, 2012), as well as in outcrop studies, including the Castagnola Formation (4–12°; Felletti, 2002; Southern et al., 2015; Marini et al., 2016); the Laga Formation (6–8°; Marini et al., 2015); and the Grès d’Annot (4–10°; Amy et al., 2007; Salles et al., 2014).

While the effects of high-gradient stepped slopes on deposition processes and architecture have been examined, the impact of subtle gradient changes on sediment distribution and flow pathways is less clear. Subtle gradient changes having a profound impact on flow processes and therefore configuration of deep-water systems have been shown from experimental (Baines, 1984; Edwards et al., 1994; Kneller and McCaffrey, 1999; Kneller and Buckee, 2000; Nasr-Azadani and Meiburg, 2014), outcrop (e.g., Kane et al., 2010; Patacci et al., 2014), and modern seabed data sets (e.g., Stevenson et al., 2013, 2014; e.g. Kneller, 1995). Despite this, previous outcrop studies on the effects of subtle fixed and dynamic topography are limited in their spatial and temporal extent (e.g., Smith, 2004b; Spychala et al., 2017a), and the influence of seabed topography over multiple depositional cycles remains unknown. Studies using seismic reflection data sets can show evolution over multiple stages of deposition (e.g., Beaubouef and Abreu, 2006; Mayall et al., 2010; Hay, 2012; Jobe et al., 2017), but they generally focus on complicated topographic templates, such as tortuous corridors (Smith, 2004a; Hay, 2012), mini-basins (Booth et al., 2003; Beaubouef and Abreu, 2006; Madof et al., 2009; Pirmez et al., 2012; Prather et al., 2012b; Oluboyo et al., 2014; Doughty-Jones et al., 2017), and piggy-back basins/foredeep margin slopes (Covault et al., 2009), and/or do not span multiple large-scale sea-level cycles (e.g., Beaubouef and Abreu, 2006; Barton, 2012). Moreover, these studies do not investigate the long-term stratigraphic evolution of slopes nor characterize stratigraphic architecture at a scale that resolves relationships between sedimentary processes and seabed topography.

This study characterizes three thin (<15 m) sandstone bodies (Units A/B, B/C, and D/E) from the Permian Laingsburg and Fort Brown formations in the Karoo Basin, South Africa. These units differ greatly in sedimentology and architecture from the larger underlying and overlying slope and basin-floor systems (Units A–F) as their accumulation did not significantly modify the slope basin-floor profile during formation. Therefore, their utility as subtle topographic indicators is investigated, building on a stratigraphic framework established from previous studies. Objectives are: (i) to document influence of seabed topography through spatial variability in thickness and facies trends; (ii) to understand the sedimentary processes and depositional environments of Units A/B, B/C, and D/E; (iii) to investigate how and when seabed topography formed and evolved, and how this topography influenced deposition of larger Units C, D, E, and F; and (iv) to develop a model for stepped slope profile evolution over multiple successive sea-level cycles. This investigation will therefore further contribute to the understanding of subtle slope evolution and its varying influence on deep-water system facies and architecture at bed to system scale.

The Karoo Basin, South Africa (Fig. 2), has been previously interpreted as a retroarc foreland basin (Visser and Prackelt, 1996; Visser, 1997; Catuneanu et al., 1998), and more recently as a thermal sag basin in the Permian that evolved into a retroarc foreland basin in the Triassic (Tankard et al., 2009). The 8-km-thick Karoo Supergroup (Fig. 2C) is subdivided into the Dwyka, Ecca, and Beaufort groups. The Dwyka Group comprises glacial deposits (Late Carboniferous to Early Permian); the Ecca Group clastic marine deposits (Permian); and the Beaufort fluvial deposits (Permian to Triassic; Belica et al., 2017). Basal deposits of the Lower Ecca Group (Fig. 2C) comprise the Prince Albert, Whitehill, and Collingham formations, mapped for 800 km along the southern margin of the Karoo Basin (Viljoen, 1992, 1994; Visser, 1992; Johnson et al., 1997). In the Laingsburg depocenter, the Collingham Formation is overlain by the Vischkuil Formation, forming the basal section of the 1800-m-thick progradational succession from basin-floor deposits (Vischkuil and Laingsburg formations; Sixsmith et al., 2004; Van der Merwe et al., 2010), through a channelized submarine slope (Fort Brown Formation; Hodgson et al., 2011; Di Celma et al., 2011; Flint et al., 2011) to shelf-edge and shelf deltas (Waterford Formation; Jones et al., 2015; Poyatos-Moré et al., 2016). Regional paleoflow is toward the northeast and east throughout the succession with the main deep-water sediment entry point located to the southwest (Van der Merwe et al., 2014). Regional-scale mapping of successive slope to basin-floor systems (Units D–F) in the Laingsburg depocenter indicates the presence of a confining, lateral basin margin to the south of the depocenter, oriented east-to-west (Van der Merwe et al., 2014; Brooks et al., 2018b).

The Laingsburg Formation comprises the stratigraphic Units A, A/B, and B, and the Fort Brown Formation, Units B/C, C, D, D/E, E, F, and G (Fig. 2C; Flint et al., 2011). This study examines Units A–F, and represents the first detailed study on the less extensive, thinner Units A/B, B/C, and D/E (Fig. 2C), that are exposed along a series of E-W–oriented subparallel post-depositional fold limbs (Fig. 2B), and intersected in a number of cored research boreholes. Detailed mapping and physical correlation of all units utilizes a regional stratigraphic framework established from previous studies (Grecula et al., 2003; Figueiredo et al., 2010, 2013; Di Celma et al., 2011; Flint et al., 2011; Hodgson et al., 2011; Kane and Hodgson, 2011; Brunt et al., 2013a, 2013b; Morris et al., 2014a, 2014b, 2016; Van der Merwe et al., 2014; Spychala et al., 2015, 2017a; Brooks et al., 2018b). The Laingsburg and Fort Brown formations have been divided into four composite sequence sets, the first comprising Unit A, the second Units A/B and B, the third Units B/C, C, and D; and the fourth Units D/E, E, and F (Flint et al., 2011). Units C, D, E, and F each represents a lowstand sequence set, with an overlying 10–30-m-thick regional hemipelagic mudstone representing the corresponding transgressive/highstand sequence set (Flint et al., 2011). Lowstand sequence sets can be further subdivided into separate depositional sequences, including a sand-rich lowstand systems tract (LST; represented stratigraphically by a subunit, e.g., E1, E2, or E3) and an overlying transgressive/highstand systems tract mudstone (∼1–5 m thick) (Figueiredo et al., 2010, 2013; Di Celma et al., 2011; Hodgson et al., 2011).

Regional mapping and correlation of units C to F have demonstrated an architectural change from sand-attached (Units C and D) to sand-detached channel-lobe transition zones (Units E and F; Van der Merwe et al., 2014; Brooks et al., 2018b). The recognition of intraslope lobes in Units D/E, E, and F (Figueiredo et al., 2010; Spychala et al., 2015), and coarse-grained bypass-dominated zones, consisting of thin channel-levee deposits and channel-lobe transition zones (sensu Mutti, 1995; Van der Merwe et al., 2014; Brooks et al., 2018b) supports the presence of a stepped slope profile at the time of E and F deposition (e.g., Figs. 1B and 1C; Van der Merwe et al., 2014; Brooks et al., 2018b). This paper focuses on the sedimentology and stratigraphic architecture of the thinner units (A/B, B/C, and D/E) in order to decipher the origin and presence of fixed and dynamic slope topography throughout the deposition of the Laingsburg and Fort Brown formations, that resulted in the formation of a stepped slope profile (Van der Merwe et al., 2014; Brooks et al., 2018b).

Over 100 stratigraphic sections were measured and logged at the mm–cm scale, to characterize the range of sedimentary facies of Units A/B, B/C, and D/E and build thickness and paleogeographic maps of each unit over tens of kilometers in depositional dip and strike. Sedimentary logs document lithology, grain size, sedimentary structures, and stratal boundaries. Beds with similar thickness, grain size, and sedimentary structures are grouped into facies. Facies are used to interpret environments of deposition within the context of extensive paleoenvironmental reconstructions already undertaken in the Laingsburg depocenter (e.g., Grecula et al., 2003; Figueiredo et al., 2010, 2013; Di Celma et al., 2011; Flint et al., 2011; Van der Merwe et al., 2014; Spychala et al., 2015, 2017a, 2017b; Brooks et al., 2018b). A physical correlation framework was established by walking out stratigraphic surfaces between sections.

Thickness distributions were created by fitting a surface to values obtained from logged sections using the kriging tool within the ArcGIS® Geostatistical Wizard. Maps are extended beyond extremities of the input data by the surfacing algorithm, with unrealistic values removed and observed trends applied. Paleogeographic maps utilize data from previous studies in the Karoo (Grecula et al., 2003; Figueiredo et al., 2010, 2013; Di Celma et al., 2011; Flint et al., 2011; Van der Merwe et al., 2014; Spychala et al., 2015, 2017a, 2017b; Brooks et al., 2018b) and represent gross depositional environments for each study interval. Unit C, D, E, and F paleogeographic maps are modified from Van der Merwe et al. (2014) incorporating data from Spychala et al. (2015) and Brooks et al. (2018b). Flow trends within the paleogeographic maps are based on restored paleocurrent data collected from ripple laminations, flutes, grooves and scours.

A 3-D datacube was constructed in Petrel® by importing 100 composite logs from outcrop and six well logs consisting of the complete Whitehill Formation to Waterford Formation stratigraphy, with the Whitehill Formation used as a basal datum. Tops and bases of key lithostratigraphic units were selected using the well top function, and regional panels were used to interpret polylines to add datapoints between outcrop logs. Depth structure maps were constructed for each key surface using the “make a surface” function and cumulative thickness maps were created between key surfaces. Post-depositional tectonic shortening was corrected by stretching 13% in the Y direction, according to a mean value for palinspastic restoration derived by Spikings et al. (2015).

Seven distinct facies are described and interpreted in terms of sedimentary processes for Units A/B, B/C, and D/E (Figs. 3 and 4).

Structureless sandstone beds are fine grained, thin to thick bedded (0.1–1 m thick) with common dish and pillar structures, and weak normal grading at bed tops (Fig. 3A). Bed geometries are tabular with erosional bases, flutes, grooves, and other tool marks (Fig. 3B). Discontinuous mudstone-clast layers are often observed at bed amalgamation surfaces, bed bases, and dispersed within beds, with sub-angular to sub-rounded clasts (<15% clasts by volume, 0.1–10 cm A-axis; Figs. 3C, 4L, and 4M). Loading and flame structures are also common at amalgamated bed contacts. Basal beds are observed to thicken and thin over meters to 100s of meters, pinching out, infilling scoured sections and onlapping or downlapping onto underlying mudstone (Figs. 5 and 6). Beds occur in packages up to 14 m in thickness, but are generally <5 m thick (Fig. 6). Bioturbation fabric (mainly Planolites) is rarely observed at bed bases.

Structureless and weakly normally graded sandstone beds suggest deposition from sand-rich high-density turbidity currents (Bouma, 1962; Lowe, 1982; Mutti, 1992; Kneller and Branney, 1995). Lack of sedimentary structures suggests rapid deposition preventing any bedform development. The presence of dispersed rip-up clasts and mudstone clast-rich amalgamated contacts suggest long-lived erosion and progressive aggradation of deposits formed by depletive steady high-density flows (Kneller and Branney, 1995). Loading and dewatering structures form syn- and post-depositionally as a result of sediment liquefaction (Mulder and Alexander, 2001; Stow and Johansson, 2000).

Structured sandstone beds are very fine to fine grained (0.05–1 m thick) with structures primarily consisting of ripple and climbing ripple cross lamination (Figs. 3F, 3H, and 4E) but also including planar lamination (Figs. 3D, 3E, and 4J) and dewatering structures (load casts and flames, and dish and pillar) (Figs. 3G, 4D, and 4F). Climbing ripple lamination can exhibit a high angle of climb (15–30°) with stoss-side preservation of laminae, which are locally sheared and overturned toward the bed top. Ripple cross-laminae can be heterolithic, comprised of both sandstone and siltstone, and mm–cm scale mud clasts (Fig. 3F). Beds occur in packages up to several meters thick (Fig. 3E). Bed geometries range from tabular to lenticular, draping surfaces as discontinuous depositional features, or as remnant erosional features that can form multiple stacked packages. Packages can stack to form thin (<3 m) and wide (>100 m) sandstone units discussed further below. Within these packages, the basal beds thicken and thin over 10s to 100s of meters, pinching out and onlapping/downlapping on to underlying mudstone unit laterally and frontally (Fig. 5). Bioturbation, mainly Planolites, is rarely observed at bed tops and bases.

Climbing ripples form from continuous bedload traction under highly depositional flows. High angle of climb, with stoss-side preservation indicates high rates of aggradation (e.g., Allen, 1970; Jobe et al., 2012; Morris et al., 2014a, 2014b), commonly associated with non-uniformity in flows and accompanying decrease in velocity and/or flow height, possibly linked to an abrupt decrease in gradient or loss of confinement (e.g., Jobe et al., 2012; Morris et al., 2014a, 2014b), driving deposition through a reduction in flow capacity (e.g., Kneller and Branney, 1995). The presence of heterolithic foresets indicates deposition from mud-rich flows (Baas et al., 2016). Millimetric to centimetric mud clasts may travel as bedload, where larger particles skip over a bed of smaller particles (Raudkivi, 1976; García, 2008); however, they might be expected to accumulate preferentially in ripple troughs. Alternatively, they may also be deposited by capacity-driven deposition from a flow containing sand and suspended mudstone clasts. Planar lamination is interpreted to have formed under upper-stage plane-bed conditions (Allen, 1984; Talling et al., 2012). Abrupt changes in basal-bed thickness and onlap terminations onto underlying mudstones are interpreted to result from an interaction with minor seabed topography with downlap representing the pinch-out of beds where flows waned sufficiently to deposit the coarser grained portion of the flow.

Beds consist of a bipartite structure with a lower division comprising weakly normally graded fine-grained sandstone (0.1–1 m thick), with some dewatering structures and rare planar lamination, and occasional mudstone-clast layers (clasts 1–10 cm a-axis). The upper division comprises poorly sorted very fine-grained sandstone and siltstone (0.1–1 m thick) with dispersed sub-angular, elongate, mm-cm scale mudstone clasts and plant fragments (Fig. 3I). Contact between divisions is abrupt and undulating (Fig. 3I).

Bipartite beds form through deposition of the lower division from a sand-rich turbidity current with the “linked” poorly sorted upper division interpreted as deposition from a co-genetic debris flow. Hybrid event beds (Haughton et al., 2003, 2009) are most commonly identified toward the bases and fringes of lobe deposits (e.g., Hodgson, 2009; Talling, 2013). However, they can form in any environment where mud and mudstone clasts are entrained into the turbulent flow, damping turbulence, and developing high-concentration to pseudo-laminar flow conditions (e.g., Ito, 2008; Haughton et al., 2003; Talling et al., 2004; Baas et al., 2011; Pierce et al., 2018).

Heterolithic siltstone and sandstone beds (<1–15 cm thick) are normally graded with planar (Fig. 4K), ripple (Fig. 4H), or climbing-ripple laminations. Beds vary in thickness laterally from meters to centimeters over 10s to 100s of meters. Basal beds are sometimes discontinuous, with pinch-out/onlap terminations onto underlying mudstones (Fig. 5). Bioturbation fabric (Planolites) is often observed, especially in siltstone dominated packages. Beds can form packages up to several meters thick (Fig. 3J), from siltstone- to sandstone-dominated (Figs. 3J and 3L), and are locally folded and deformed (Fig. 3K).

Normally graded siltstone beds suggest deposition from dilute turbidity currents, with the finer sediment residual within the flow after deposition of the coarser fraction of sediment load (Stow and Bowen, 1980; Jobe et al., 2012). Climbing-ripple lamination forms through late-stage tractional modification by flows, with high-sediment fall-out rates (Lowe, 1988). Low angle of climb suggests higher migration component and lower rates of suspended load fallout and aggradation than in structured sandstone (Allen, 1970; Jobe et al., 2012). Localized deformation and folding indicates slumping of material and/or significant dewatering. Thickness variations, discontinuities and onlap of basal beds are interpreted to result from deposition on irregular seabed topography.

Chaotic beds comprise contorted sand clasts supported by a poorly sorted silt-prone matrix (Fig. 3M). A variant of this facies are poorly sorted, matrix-rich, very fine-grained sandstone to coarse siltstone beds (0.1–2 m thick) that lack internal structure and contain dispersed sub-angular, elongate, mm-cm scale mudstone clasts and plant fragments (Figs. 3M and 4B).

The beds with clasts supported by a poorly sorted matrix are interpreted to have been deposited en masse, and named as debrites (debris-flow deposits; Mulder and Alexander, 2001; Talling et al., 2012).

Units of contorted and folded strata comprising thin- to thick-bedded sandstones and siltstones can be up to 4 m thick and extend laterally for 10s of meters (Fig. 3N).

These units are interpreted as remobilized deposits formed by slides and slumps.

Claystones and siltstones that have mm-scale laminations are structureless (Fig. 4). This facies forms regionally extensive units that have been mapped for 10s of km (Fig. 3O) and drape every sand-prone unit (Flint et al., 2011; Van der Merwe et al., 2014).

These deposits are interpreted as background hemipelagic deposition.

The thin sandstone Units A/B, B/C, and D/E are interpreted as deep-water lobes due to lateral continuity of beds, lack of major erosion surfaces, and their constituent facies. Four broad environments of deposition have been characterized (Fig. 6) based on the grouping of facies outlined above, mapped geometries, and paleogeographic context, utilizing depositional environment interpretations from previous studies (Prélat et al., 2009; Kane and Hodgson, 2011; Prélat and Hodgson, 2013; Brunt et al., 2013a, 2013b; Morris et al., 2014a, 2014b; Spychala et al., 2015).

Lobe axis: Consisting primarily of thick amalgamated structureless sandstone facies (commonly containing mudstone-clast horizons) interbedded with minor structured sandstone facies (Fig. 6). Lobe axis packages are up to 14 m thick. Typically, contacts with the underlying mudstone are sharp and sometimes scoured with numerous flutes, grooves and other tool marks. In proximal areas of the depocenter, deposits comprise structureless and structured sandstone facies as well as thin-bedded sandstone and siltstone facies. Structured sandstone facies form thin (<5 m) and wide (>100 m) lens shaped packages (Figs. 5E and 6) that onlap, drape, and overtop underlying seabed topography (Fig. 5E). Indicators of sediment bypass include multiple composite scours, with amalgamated erosional surfaces (cm–1.5 m in depth, 0.5–3 m in width) mantled with mudstone-clast lag horizons (e.g., Hofstra et al., 2015; Stevenson et al., 2015).

Lobe off-axis: Consisting of stratified structureless and structured sandstone facies and sand-dominated thin bed facies (Fig. 6). Lobe off-axis packages are up to 5 m thick. Scouring is documented at bed bases and minor erosion is occasionally present throughout. Packages often thicken and thin abruptly with bed thickness changes and onlap of the lower beds onto regional mudstones, due to compensation and infilling of topographic lows.

Lobe fringe: Frontal and lateral lobe-fringe deposits consist of thin structured/structureless sandstone facies with common mud-clast horizons, thin-bedded sandstone and siltstone facies, bipartite beds, and debrites (Fig. 6). Beds are mm to 50 cm thick with lobe fringe packages of thin beds up to 2.5 m and sandstone up to 70 cm. Sand-rich pinch-outs are often associated with clastic injectite networks, up to several meters in thickness and kilometers in lateral extent (Cobain et al., 2017).

Distal lobe fringe: Where fringes are silt-rich, at distal pinch-outs (Fig. 6), deposits consist of thinly bedded siltstone facies with occasional starved ripples. These are distinguished from the regional mudstones as they are coarser than the background sediments and have been walked out from more proximal areas. It may be possible that other packages of distal lobe-fringe facies are present but undocumented.

Hemipelagic drape: This consists of regional mudstone facies and represents periods of clastic starvation in the basin when hemipelagic material can accumulate to regionally extensive, significant (2–10s of meters) thicknesses.

Remobilized packages: These result from syn-sedimentary deformation and can be associated with all environments of deposition, including lobe fringe (Spychala et al., 2017a); levee collapse (Kane and Hodgson, 2011; Morris et al., 2014a); channel margin (Hodgson et al., 2011) and hemipelagic drape (Brunt et al., 2013b). Deposits may represent localized remobilization of packages due to instabilities caused by minor seabed topography (e.g., Spychala et al., 2017a).

The thin sandstone units are present discontinuously in distinct, spatial areas across the Laingsburg depocenter, with deposition generally concentrated along the western, up-dip Baviaans and Heuningberg locations (Fig. 7). Typically, Units A/B, B/C, and D/E are an order of magnitude thinner than Units A–F (generally <5 m in thickness; Fig. 6), reaching a maximum of 14 m in thickness (Fig. 7). Overall, the thicker areas are dominated by the lobe axis deposits, transitioning to off-axis and fringe deposits laterally and frontally as the unit thins (Figs. 6 and 7).

In the Baviaans area, flute and groove paleocurrents are consistently toward the north-northeast, changing to eastward around Heuningberg (Fig. 7). Current ripple paleocurrents are more dispersed showing a spread between north-northwest and east. The thickest A/B deposits (up to 14 m) are aligned in a north to south trend, coinciding with the most axial lobe facies (Fig. 6). The unit thins (14 m–1 m) and transitions to off-axis deposits more abruptly (over <3 km) toward Faberskraal and Geelbek to the east and more gradually (over 20 km) to the west, with fringe deposits restricted to the peripheries of the system in Heuningberg, Faberskraal, and Geelbek (Fig. 7). A localized area of increased thickness (up to 10 m) around Heuningberg, thins in all directions (Fig. 7). Lateral pinch-outs (Faberskraal, Geelbek, and Heuningberg) are indicated by a combination of highly discontinuous sand-rich strata and silt-rich thin beds (Figs. 6 and 7). In the most proximal Baviaans area, high aspect ratio erosional features (1–5 m in depth and 10s of meters to 150 m in width) are present over a 10 km strike-oriented outcrop section (Figs. 5E, 6, and 7). Erosional surfaces are overlain by structureless, planar and climbing-ripple laminated sandstone beds that thin and fine laterally to interbedded siltstone and ripple-laminated sandstone (Figs. 5E and 6).

Unit B/C is also best developed in the regional up-dip areas of the depocenter (Baviaans, Heuningberg, and Geelbek; Fig. 7) with thickest and most axial deposits around Baviaans. Flute, groove, and ripple paleocurrents indicate a northeast direction of transport, with a northwest component at Geelbek (Fig. 7). Unit B/C is thinner overall than A/B (<4 m thick) and less laterally extensive (Fig. 7). Toward Heuningberg and Faberskraal, deposits transition to off-axis and pinch out more abruptly (Fig. 7). A discontinuous area of thicker (up to 4 m), sand-rich lobe-fringe deposits is present around Geelbek, thinning more abruptly to the west, north, and east, likely related to unexposed axial deposits to the south (Fig. 7). Pinch-outs are sand-rich at the lateral (eastern) fringe (Fig. 6) and more silt-rich and thin-bedded (Fig. 6) toward the frontal (northern) fringe (Fig. 7).

Unit D/E has been recognized in three discrete sites, correlated though walking key regional markers including top Unit D, base Unit E, and the D-E mudstone. Depositional sites are around Heuningberg, Geelbek, and Floriskraal (Fig. 7). In Heuningberg, up to 14 m of axis and off-axis lobe deposits are present, that thin and transition to fringe deposits to the west, south, and east, pinching out more abruptly to the south and east and more gradually to the west (Fig. 7). In Geelbek, Unit D/E consists of a single debrite bed <2 m thick (Fig. 7). In Floriskraal, a 14-m-thick lobe axis has a southwest to northeast depositional dip oriented pinch-out with paleoflow to the northeast and east (Fig. 7). Unit D/E is highly discontinuous toward Slagtersfontein, transitioning to off-axis and fringe deposits, and pinching out more abruptly to the south and more gradually to the north (Fig. 7). Pinch-outs are more sand-rich (Fig. 6) in the frontal (eastern) fringe and more silt-rich and thin-bedded (Fig. 6) in the lateral (southern) fringe with remobilized/debritic deposits present in the northern lateral fringe (Fig. 7).

Each unit is bounded by a regional mudstone. Locally, below the thin sandstone units silt-rich remobilized strata are present, centimeter to several meters in thickness, ranging from chaotic and highly disaggregated to laterally extensive (meters to 10s of meters).

Facies and thickness distributions indicate A/B and B/C systems were fed from the southwest (Figs. 7 and 8), with mean paleoflow toward the northeast. Lenticular erosive features in the Baviaans area may represent either: shallow scours, often associated with channel-lobe transition zones (e.g., Wynn et al., 2002; Macdonald et al., 2011a, 2011b; Hofstra et al., 2015; Pemberton et al., 2016; Brooks et al., 2018b) formed by hydraulic jumps where flows transition from super- to subcritical due to a reduction in slope gradient and/or flow confinement (Mutti and Normark, 1987, 1991; Weirich, 1989; Kostic and Parker, 2006; Sumner et al., 2013; Dorrell et al., 2016); or weakly confined distributive channels eroding into proximal lobes, with distributive patterns likely due to the reduction of slope gradient (e.g., van der Werff and Johnson, 2003). These characteristics indicate deposition in a base-of-slope area. The distribution of Unit D/E indicates at least two areas of sediment input, in the northeast and to the south (Figs. 7 and 8).

Similar geographical areas (Faberskraal and Geelbek) of deposition, directions of facies transitions, and thinning and pinch-outs are recognized in all three units (Figs. 7 and 8). Thinning and pinch out of A/B and B/C lobes occur in Heuningberg, where a southeast-facing intrabasinal slope is interpreted to impact flow behavior in the underlying Unit A (Spychala et al., 2017a). We suggest that this intrabasinal slope was present throughout A/B and B/C deposition (Figs. 7 and 8), modifying facies, thickness and paleocurrent trends by reflecting and deflecting flows. The influence of this intrabasinal slope does not persist into Unit D/E deposition indicating that topography was healed by intervening units (Fig. 8).

Eastern pinch-outs in Units A/B, B/C, and of D/E consistently occur across the Faberskraal-Geelbek area (Figs. 7 and 8), suggesting influence by some form of long-lived and relatively fixed, topographic feature (Fig. 8). Lobe pinch-outs indicate a decrease (or slight reversal) in slope gradient, confining deposition up-dip. This topography was apparently fixed in location, but evidently subtle or dynamic, as the effect of large-scale topographic depression is not recognized in the intervening units (Van der Merwe et al., 2014). Therefore, this topography must have formed progressively over time, with a subtle expression on the seabed at any one time (e.g., Spychala et al., 2017a).

The abrupt southward pinch-out of Unit D/E in Heuningberg (Fig. 7) may relate to a north-facing slope, resulting from differential compaction over underlying stratigraphy (Figueiredo et al., 2010). The southwest to northeast–oriented deposition of the Unit D/E axis in Floriskraal (Fig. 8) and pinch-out at Slagtersfontein may be the result of a southeast-facing slope surface representing the regional base-of-slope (Fig. 8). The southward thinning and pinch-out is likely the result of a north-facing lateral basin margin (Fig. 8) (Van der Merwe et al., 2014; Brooks et al., 2018b).

Remobilization within regional mudstones prior to A/B, B/C, and D/E deposition is interpreted to indicate instability of locally steepened frontal and lateral slopes that subsequently controlled lobe deposits and potentially created additional local topography.

The dominance of laterally extensive, unconfined, basin-floor lobe deposits within Units A and B (Sixsmith et al., 2004; Prélat and Hodgson, 2013; Van der Merwe et al., 2014) indicates Unit A/B was deposited in a similar basin-floor setting (Fig. 8). Units C, D, and E are interpreted as slope to basin-floor systems (Figueiredo et al., 2010; Van der Merwe et al., 2014; Morris et al., 2014a, 2014b, 2016; Spychala et al., 2015; Brooks et al., 2018b), with the western, up-dip sections on the slope. The co-location of Unit B/C in this area indicates a lower slope setting. Unit D/E in Heuningberg and Geelbek likely represents intraslope lobes (Spychala et al., 2015), with the Floriskraal deposits on the basin floor (Fig. 8).

Unit A/B, B/C, and D/E lobe deposits are subdivided into axis, off-axis and lateral/frontal fringe (Fig. 6), with similar facies transitions in lobes and lobe complexes (sensu Prélat et al., 2009) recognized in the larger units. These units rarely record juxtaposition of multiple sub-environments (Fig. 6), indicating that deposition was not sufficiently sustained or sediment volume was not adequate for lobes to stack. Neither compensational stacking to form a lobe complex, nor propagation of channel-levee systems above lobes (e.g., Morris et al., 2014b; Hodgson et al., 2016) is documented in these units. In addition to thickness, A/B, B/C, and D/E show key differences in facies and architecture to the larger Units A–F (cf. Sixsmith et al., 2004; Prélat et al., 2009, 2010; Prélat and Hodgson, 2013). These include (a) a higher proportion of sand (Figs. 4 and 6) particularly in frontal fringes (Fig. 6); (b) consistently sharp bases and tops (Figs. 4 and 6); (c) a higher proportion of mudstone clasts (Figs. 4 and 6); (d) minor (<1 m) basal scouring throughout, including at fringes (Figs. 5 and 6); (e) significant and widespread abrupt thinning and pinch out causing discontinuity (Figs. 5 and 7), rarely associated with facies changes; and (f) weakly confined channels/scours in proximal area of lobes. Similar to the larger units, lateral fringes are recognized as more silt-rich and thin-bedded whereas frontal fringes are sand-rich and discontinuous (Fig. 6). This variation is also documented in the Tanqua depocenter (Rozman, 2000; Prélat et al., 2009; Groenenberg et al., 2010) and is interpreted to relate to differences in flow processes in distal frontal versus distal lateral run outs (Spychala et al., 2017b).

Units A/B, B/C, and D/E stratigraphically overlie the thickest regional mudstone units in the Laingsburg and Fort Brown formations (Grecula et al., 2003; Di Celma et al., 2011; Van der Merwe et al., 2014), and are interpreted as the deposits from the first flows after the longest period of clastic shut off in the deep water part of the basin, reflecting the longest duration relative sea-level rise (Flint et al., 2011). Flint et al. (2011) suggested these units were distal expressions of larger scale lobe complexes similar to Units C–F. However, key differences in sedimentology and architecture described above indicate this is not the case. The absence of persistent supply that may drive lobe stacking and propagation of channel-levee systems (Fig. 8) suggests the thin units may not have been fully connected to upper slope and shelf feeder systems (Fig. 9). Delta propagation to the shelf edge, or across-shelf sediment bypass, may have been insufficient to activate upper slope (e.g., canyon) systems and sediment may have been sourced directly from the outer shelf region and via slope failures (Fig. 9). A range of input directions may have resulted in weakly confined flows, unrestricted to long-lived conduits (e.g., Saller et al., 2004; Moody et al., 2012; Stevenson et al., 2013). Multiple, short-lived sediment input locations would create widespread areas of lobe deposition. The interaction of deposits produced by these multiple inputs, would lead to lobes that do not stack compensationally or aggradationally (Fig. 9). A lack of significant channelization could cause flows to traverse completely “out-of-grade” areas (sensu Prather, 2003) that would erode the slope, driving entrainment of mud clasts. Unconfined flows would spread laterally (Al Ja’Aidi et al., 2004; Kane et al., 2008; Hamilton et al., 2017) and be more susceptible to topographic perturbations than more channelized flows (Kneller and Buckee, 2000).

Disconnected upper slope feeder systems would suggest that relative sea-level falls were not of sufficient magnitude or duration to expose the shelf or bring the deltas close to the shelf edge (Fig. 9). Therefore, small-scale rises in relative sea level (or autogenic adjustment of the shelf delta or slope equilibrium) would be sufficient to cut off sediment supply to the deep water. Initiation of Units A–F is interpreted as marking times when relative sea-level fall was sufficient to expose the shelf and/or bring the delivery systems to the shelf edge and activate upper slope feeder systems (Fig. 9). The lobe deposits of Units A/B, B/C, and D/E are therefore interpreted as “disconnected lobes” that are characterized by supply from short-lived conduits, such as slope gullies or locally developed incipient channels, rather than point-sourced flows through mature channel-levee systems (Fig. 9). It is also possible that other methods of long-term (i.e., 100,000s to millions of years’ time scale) allogenic forcing could have influenced shelf and slope sediment delivery systems resulting in the variation of sediment input.

These disconnected lobes can be compared to models of intraslope lobes (also known as perched lobes; Plink-Björklund and Steel [2002]; Prather et al. [2012a]; or transient fans; Adeogba et al. [2005]; Gamberi and Rovere [2011]). These models based on shallow seismic reflection data sets have been interpreted to show initial stages of fully confined ponded facies, transitioning stratigraphically through partially confined wedge shaped fill, and finally bypass stages (Prather et al., 1998; Prather, 2000). In contrast, the disconnected lobes herein show no evidence for initial ponding, although onlap and downlap is documented. The absence of graded bed tops throughout the units and lack of fines suggest downstream flow-stripping (Sinclair and Tomasso, 2002; Spychala et al., 2015) occurred, and that the upper, finer portions of the flow were able to surmount the topographic confinement and bypass down-dip. This indicates slope topography was not of sufficient magnitude or three-dimensional constraint to completely pond flows (e.g., Haughton, 1994). Moreover, although units have a sharp top, late-stage sediment bypass is also not documented in these disconnected lobes suggesting slope accommodation remained underfilled.

The depositional trends documented in Units A/B, B/C, and D/E indicate subtle seabed topography was maintained over relatively long periods, resulting in a persistent location of pinch-out. Here, we discuss topographic indicators present in the thicker, more laterally extensive Units C, D, E, and F (Figs. 10, 11, 12, and 13A), that have been mapped in detail in previous studies (Grecula et al., 2003; Figueiredo et al., 2010; Di Celma et al., 2011; Hodgson et al., 2011; Brunt et al., 2013a; Van der Merwe et al., 2014; Spychala et al., 2015; Morris et al., 2016; Brooks et al., 2018b).

Intraslope lobe complexes have been recognized in subunits E1, E2, and E3 with maximum thickness of 5, 15, and 20 m, respectively, and within subunits F1 and F3 with maximum thicknesses of 5 and 35 m, respectively (Fig. 11) (Figueiredo et al., 2010; Van der Merwe et al., 2014; Spychala et al., 2015). All intraslope lobe complexes have an eastern down-dip pinch-out in the Faberskraal-Geelbek area (Figs. 11 and 13A), marking the presence of a persistent area of local accommodation (Fig. 11) (Figueiredo et al., 2010; Van der Merwe et al., 2014; Spychala et al., 2015). At any one time, the amount of accommodation would have been subtle, but long-term development in fixed locations led to a marked impact on the stratigraphic architecture of the basin fill (Figs. 12 and 13A).

Unlike the smaller fan units, these intraslope lobe complexes (Units E and F) demonstrate a stratigraphic transition through initial deposition and infill of slope accommodation by lobe deposits (E2 and F2; Figure 11) and later stage sediment bypass through channel-levee systems (E3 and F3; Figure 11), similar to that documented in the Niger Delta slope (Prather et al., 2012a; Jobe et al., 2017; Fig. 1E) and various basins throughout the Gulf of Mexico (Prather et al., 1998, 2017; Fig. 1D). This indicates that sediment input was sufficient for slope accommodation to become healed.

The long-lived feeder system present around Baviaans throughout Units C and D time has been well documented (Fig. 10) (Hodgson et al., 2011; Kane and Hodgson, 2011; Brunt et al., 2013a; Morris et al., 2014a, 2016). This system varied from entrenched up-dip to more distributive down-dip, and has been documented within subunits C2, D1, D2, and D3 across the Faberskraal-Geelbek area (Fig. 10) (Brunt et al., 2013a; Van der Merwe et al., 2014; Morris et al., 2014a, 2016). Due to the spatial coincidence with intraslope lobe complex pinch-outs (Figs. 10 and 13A), it is possible that a reduction in slope gradient caused flows to decelerate, decreasing their ability to erode.

Zones of coarse sediment bypass (sensu Stevenson et al., 2015) have been recognized within Units E and F (Van der Merwe et al., 2014; Brooks et al., 2018b), within the central area of the depocenter, between Faberskraal/Geelbek and N1 Dome/Slagtersfontein (Figs. 11 and 12), indicating a persistent area of decreased accommodation, on a slope close to equilibrium (Prather, 2003). The zone is characterized by thin channel-levee deposits and the formation of sediment bypass-dominated channel-lobe transition zones (CLTZs) (e.g. Mutti 1985; Mutti and Normark, 1987, 1991; Wynn et al., 2002) at Slagtersfontein within Unit E (Brooks et al., 2018b) and the N1 Dome within Unit F (Figs. 11 and 12). These CLTZs represent the regional base-of-slope indicated by paleogeographic context and facies transitions from channel-levee systems to unconfined basin-floor lobe complexes (Brooks et al., 2018b; Fig. 11). This slope is interpreted as southeast facing during Unit E deposition, due to the orientation of up-dip lobe pinch-outs (Fig. 11). Zones of coarse-sediment bypass in Units E and F likely indicate an increase in slope gradient in this central area in order for flows to bypass efficiently and the development of the step-ramp-step geometry (Brooks et al., 2018b; Fig. 11). Absence of Units D/E and subunit D3, as well as the thinning of subunit D2 in the central area, may be early indicators of this ramp slope (Figs. 7, 10, and 12).

Within subunits E2 and E3, tabular, laterally continuous (kilometers in dip and strike), thin bedded siltstone packages (spill-over fringe deposits) are recorded beneath the basin-floor lobes (Brooks et al., 2018b; Fig. 11), and reflect initial flow stripping of fines and trapping of sand in intraslope lobes (Sinclair and Tomasso, 2002). Later channels incised into the intraslope lobes (Fig. 11) (Spychala et al., 2015), indicating that accommodation was filled and the slope was at grade.

In order to understand when the fixed seabed topography initiated, a 3-D datacube was constructed in Petrel® using the laterally persistent Whitehill Formation as a basal datum and regional mudstones to correlate units throughout the entire Laingsburg depocenter. Figure 13B shows a dip section of the Vischkuil, Laingsburg, and Fort Brown formations, highlighting two distinct areas of increased thickness separated by a zone of thinning in all units. Thickness maps have also been constructed using this datacube, from the early stage starved basin plain deposits and mass-transport deposits of the Collingham and Vischkuil formations through the overlying basin-floor lobes of Units A and B in the Laingsburg Formation (Fig. 13B). The Collingham, Vischkuil, and Laingsburg formations (Fig. 13B) have the same general thickness patterns as the overlying units, with thicker deposits in the west, thinning in the central area and thickening to the east and southeast.

The repeated thickness and facies trends across all units (Figs. 12 and 13) indicate the presence of fixed locations of persistent topographic influence and suggest that the systems were not “at grade” (sensu Prather, 2003) at the initiation of deposition. An “out of grade” slope was generated during deposition of the thick regional mudstones of the transgressive and highstand sequence sets within each composite sequence. The rate of accommodation creation was outpaced by sedimentation during deposition of lowstand clastic units. Therefore, longer hiatuses in sand input, inferred from higher mudstone thicknesses (see Figs. 12 and 14), would lead to higher amplitude intraslope accommodation, with Units A/B, B/C, and D/E deposited during these times.

The thinning of the Vischkuil Formation as well as Units A and B (Fig. 13B), the absence of A/B, B/C, and D/E (Fig. 7), the thinning of Units C and D (Fig. 10), and the formation of sediment bypass dominated zones in Units E and F (Fig. 11) all occur in the central area (Figs. 12 and 13). This indicates a long-lived area of reduced accommodation, interpreted as a long-lived higher gradient “ramp” within a stepped slope profile (e.g., Figs. 1B and 1C). This stepped profile had a more acute impact in the later Units E and F (Figs. 11 and 12) (Van der Merwe et al., 2014; Brooks et al., 2018b), but was present to some extent from the onset of deep-water sand supply to the depocenter (Fig. 13B).

The cause of the protracted deformation of the slope is unknown, but it is clear that it was active prior to the onset of sand deposition (Unit A) in the basin (Fig. 13B). Influence of inherited seabed topography alone (e.g., Adeogba et al., 2005; Gamberi and Rovere, 2011; Olafiranye et al., 2013) cannot be the reason as it would be healed over time. Dynamic mechanisms such as salt diapirism and mobile shale (e.g., Barton, 2012; Deptuck et al., 2012; Hay, 2012; Prather et al., 2012a, 2012b), have the potential to create significant and reoccurring accommodation such as in the Gulf of Mexico (Prather et al., 1998; Prather, 2000; Meckel et al., 2002), but are not present in the stratigraphy. Active tectonic structures can also create topographically complex slopes (e.g., Hodgson and Haughton, 2004; Burgreen and Graham, 2014). However, no large-scale syn-sedimentary tectonic structures are documented in the Karoo Basin, although it is possible that structures could be present at high angles to the post-depositional fold structures associated with the deformation front of the Cape Fold Belt. The characterization of the Karoo Basin as a thermal sag basin related to subduction dynamic topography (Tankard et al., 2009) suggests that differential subsidence or localized uplift could be possible mechanisms for generating this deformation. The repeated accommodation created west of the Faberskraal-Geelbek area indicates this area was undergoing subsidence at an increased rate compared to east of the area. This could indicate increased subsidence relative to the overall subsiding basin, or a stationary/uplifting basement block east of Faberskraal-Heuningberg, resulting in a steeper slope, leading to the formation of the E and F sediment bypass dominated zones. Thermal subsidence can vary with upper crustal heterogeneity or thickness of basement, which are variable across the Karoo Basin (Binks and Fairhead, 1992; Bumby and Guiraud, 2005) or with spatial variations in mantle downwelling and crustal heterogeneity (Tankard et al., 2009), with magnetic anomalies in the underlying basement (Weckmann et al., 2007; Tankard et al., 2009; Lindeque et al., 2011) indicating the existence of a highly complex upper crust.

Overall, three key stages of stepped slope evolution are interpreted in the Laingsburg depocenter (Fig. 14). In the absence of a strong chronostratigraphic framework, the rate of slope deformation is assumed to be constant, with variations in sediment input and the stacking of systems resulting in varying stages of topographic influence.

• Phase 1 (Fig. 14; Units A, A/B, and B). These units were deposited at a time of high sediment supply, outpacing deformation rate, healing accommodation created during times of sand shut off (Fig. 14). During early stages of Unit A sedimentation, healed slope accommodation (sensu Prather, 2000, 2003) dominated. Due to the longer term sand shut off between the Unit A and B composite sequence sets, it is likely Unit A/B and initial Unit B flows were deposited in partially confined healed slope accommodation. During Unit A and B deposition, sedimentation rate was sufficient to outpace the deformation rate, allowing depositional and erosional patterns to modify the basin-floor profile toward equilibrium.

• Phase 2 (Fig. 14; Units B/C, C, and D). These units were deposited at a time of moderate sediment supply, in balance with deformation rate, with phases of healing and degradation of the slope (Fig. 14). The shut-off of sand supply during the B-B/C transgressive and highstand sequence set caused B/C to commence deposition on a deformed slope (Fig. 14). Therefore, Unit B/C and initial Unit C sedimentation likely occurred above partially confined slope accommodation with continued C and D deposition filling and healing slope accommodation (sensu Prather, 2003). The entrenched channel systems (Fig. 10) show that slope profile was above grade and undergoing erosion to establish a graded profile (e.g., Pirmez et al., 2000; Deptuck et al., 2012; Hay, 2012).

• Phase 3 (Fig. 14; Units D/E, E, and F). These units were deposited at a time of lower sediment supply, with deposition commencing on a more deformed slope after the D-D/E transgressive and highstand sequence set sand shut-off. Units D/E, E1, F1, and F2 were deposited in intraslope accommodation, with E2, E3, and F3 deposited in healed slope accommodation (sensu Prather, 2003) where they evolved to “fill and overspill” accommodation to bypass sediment down-dip (e.g., Prather et al., 1998; Beaubouef and Friedmann, 2000; Pirmez et al., 2000; Booth et al., 2003; Barton, 2012; Bohn et al., 2012).

Overall, this evolution demonstrates the transition from a relatively “simple” slope to basin-floor profile (Fig. 1A) to a stepped slope profile (Figs. 1B and 1C) with development of healed, but not ponded accommodation (sensu Prather, 2003). Previous studies have demonstrated that stepped slopes evolve temporally, often attributed to active mobile substrates or tectonics (Prather, 2003; Hay, 2012; Jobe et al., 2017; Fig. 1E), healing of mini-basins (Satterfield and Behrens, 1990; Prather et al., 1998; Prather, 2003), or through healing of initial mass-wasting generated slope relief (Shultz and Hubbard, 2005; Romans et al., 2011; Brooks et al., 2018a). An example of topographic relief being healed through time (Phase 1, Fig. 14) is the Tres Pasos Formation, Chile, where initial mass wasting resulted in an out of grade slope profile (Shultz and Hubbard, 2005; Armitage et al., 2009) that was subsequently smoothed to permit progradation of the slope (Hubbard et al. 2010; Romans et al., 2011). However, exhumed submarine slope systems with the required stratigraphic control and spatial extent to document increasing topography of submarine slope systems are rare and tend to come from small tectonically active basins with steep slopes (e.g., Haughton, 1994; Haughton, 2000; Hodgson and Haughton, 2004; Kane et al., 2010; Moody et al., 2012). Therefore, this study of the impact on stratigraphic architecture during long-term evolution of subtle slope topography over a large area is unique.

This study demonstrates that stepped slope profiles (e.g., Figs. 1B and 1C) can evolve from relatively “simple” slope profiles (e.g., Fig. 1A) without mobile substrate (e.g., Figs. 1D and 1E), significant tectonic deformation or mass wasting, instead forming through subtle slope deformation, stacking of multiple systems and significant periods of coarse clastic shut off, allowing deformation to outpace depositional healing (Fig. 14). This highlights a key consideration suggested by previous seismic reflection studies (e.g., Prather, 2003; Jackson et al., 2008) that the size and scale of deformation compared to sedimentation rate is critical, and that systems with episodic sediment flux form more slope accommodation than those without (Prather, 2000, 2003). We have shown that subtle, but slow, and persistent deformation over a significant area of a slope (many kilometers) with significant periods of sand shut off (e.g., between Units D and E deposition) can modify sediment delivery pathways and consequently sand distribution to an extent comparable to that generated by more obvious mobile substrate and tectonics. This has implications for hydrocarbon reservoir prediction within slope systems (e.g., Western Niger Delta, Deptuck et al., 2012; Gulf of Mexico, Prather et al., 1998, Prather, 2000, Meckel et al., 2002; and Offshore Angola, Hay, 2012) which have become increasingly significant as exploration targets in many basins worldwide (e.g., Weimer and Link, 1991; Mayall et al., 2006).

Thickness and facies distributions of Units A/B, B/C, and D/E, including persistent areas with the presence or absence of sand deposition, indicate a strong influence of fixed seabed topography on the slope and basin floor. These units vary from the larger units in the Laingsburg and Fort Brown formations, due to their thin, sand-rich nature, with evidence of increased erosion, and lack of lobe stacking or channel propagation. These thin units are interpreted to represent the first sand supply to the deep basin following extended periods of high sea levels. Characteristics of the thin units are interpreted as marking only partial re-establishment of sand delivery pathways, here termed “disconnected lobes,” which are not sourced from major feeder channel-levee systems, but from periodic failures on the shelf edge and upper slope. Such disconnected lobes vary from intraslope lobes in the Laingsburg depocenter and other systems, due to their stratigraphic position and limited deposit thickness, as a result of their formation within the initial stages of long-term relative sea-level fall. In contrast, larger Units A–F likely formed in periods of lower sea level that were sufficient to bring delivery systems to, or close to the shelf edge, developing and reusing mature upper slope sediment distribution pathways.

Persistent areas of increased or decreased thickness in the underlying formations suggest that the topography was developed before onset of major sand input to the basin. Overall, three stages of evolution are recognized, with a temporally increasing impact on slope to basin-floor systems. In Phase 1, sedimentation rate outpaced deformation rate. During Phase 2, sedimentation rate was roughly equal to deformation rate. Finally, in Phase 3, sedimentation rate was outpaced by deformation rate, with the stepped slope topography showing a dominant control on depositional architecture. The cause of deformation is unknown but may relate to differential subsidence across basement heterogeneities. The model presented here can be applied to identify accommodation change and deep-water system architecture across evolving stepped slope profiles. This study indicates that subtle but persistent deformation, paired with significant periods of clastic shut off, can have a comparable impact to mobile substrate, active tectonics, or mass wasting on sediment routing.

Flows sourcing the thinner discontinuous sand units tend to deposit in topographic lows and deflect around highs, consequently persistent areas of presence/absence can indicate long term topographic influence, making these units more reliable topographic indicators than larger units which often modify the slope profile as they deposit.

The authors thank the local farmers of the Laingsburg area for permission to undertake field studies on their land. De Ville Wickens helped with regional context and logistical support. We thank Rachel Harding for the construction of datacube figures. We thank Sarah Cobain, Florian Pohl, Lewis Burden, Sophie Cullis, and Grace Cosgrove for field assistance. This work was carried out as part of the SLOPE 4 consortium research project. We are grateful for financial support from: Anadarko, BHP Billiton, BP, ConocoPhillips, ENGIE, Maersk Oil, Murphy, Nexen, Petrobras, Premier Oil, Shell, Statoil, Total, VNG Norge, and Woodside. The authors would also like to thank Bradford Prather, Matthew Malkowski, Associate Editor Andrea Fildani, and Science Editor Shanaka de Silva for their insightful comments and helpful suggestions to improve the manuscript, and Zobe Jobe and Bradford Prather for use of original graphics adapted in Figure 1.