The succession in the Chacay Melehue area was deposited in a half-graben (∼20 km long) (Manceda and Figueroa, 1995; Llambías et al., 2007; Leanza et al., 2013) that occupied the western and deepest part of a broader early post-rift depocentre in the Central Neuquén Basin (∼65 km long) (Manceda and Figueroa, 1995; Veiga et al., 2013). The half-graben shows a strong asymmetry due to a steep western margin characterised by the development of the early Andean magmatic arc and location of a major syn-rift fault (Manceda and Figueroa, 1995; Suárez and de la Cruz, 1997; Vicente, 2005), which contrasts with the stable and gently dipping eastern cratonic margin (Spalletti et al., 2012; Veiga et al., 2013). Deposition of the Los Molles Formation took place during a period of thermal subsidence and regional marine transgression across complex inherited rift topography, which promoted the reduction of sediment supply and sand starvation in this part of the basin (Spalletti et al., 2012; Veiga et al., 2013). The proximity to the volcanic arc (∼30 km to the west), the abundant volcaniclastic deposits (Zöllner and Amos, 1973; Rosenfeld and Volldaeimer, 1980; Gulisano and Gutiérrez Pleimling, 1995; Suárez and de la Cruz, 1997; Vicente, 2005; Llambías et al., 2007), and palaeocurrent measurements in sandstones indicating a general southeastwards trend reveals that sediment supply feeding the Chacay Melehue area during the post-rift was sourced from the western magmatic arc (Gulisano et al., 1984; Vicente, 2005). The deep-marine deposits of Los Molles Formation (Weaver, 1931) overlie shallow-marine tuffaceous clastic deposits (De La Cruz and Suarez, 1997; Llambías and Leanza, 2005) and carbonate deposits of the Chachil Formation (Pliensbachian to Early Weaver, 1942; Kamo and Riccardi, 2009; Leanza et al., 2013; Riccardi and Kamo, 2014), deposited with the first marine incursion in the basin (Gulisano and Gutiérrez Pleimling, 1995; Leanza et al., 2013) (Figure 2A). Chronostratigraphic studies based on ammonite biostratigraphy (Gulisano and Gutiérrez Pleimling, 1995; Riccardi, 2008) and U-Pb radiometric dating (Kamo and Riccardi, 2009; Leanza et al., 2013; Riccardi and Kamo, 2014), place the Los Molles Formation in the Chacay Melehue region as Early Toarcian-to-Early Callovian in age (Gulisano and Gutiérrez Pleimling, 1995) (Figure 2C). The succession of the Los Molles Formation in the Chacay Melehue depocentre is 850 m thick (Figure 2B). A 55 m thick sandstone-prone interval in the lower succession represents an Aalenian turbidite system (interval II of Gulisano and Gutiérrez Pleimling, 1995). The overlying Bathonian section of the Los Molles Formation (up to 200 m thick) (Figure 2B) is mainly represented by mudstone and heterolithic successions, including a 70 m thick interval (study interval; Figure 2C) of deformed sand- and mud-rich deposits (interval IV of Gulisano and Gutiérrez Pleimling, 1995). The overlying Lower Callovian strata of the Los Molles Formation is characterised by a 300 m thick interval of thin-bedded mudstone. It is overlain by either the fluvial Lotena Formation (Gulisano and Gutiérrez Pleimling, 1995; Veiga et al., 2011) or evaporites (Tábanos Formation; Figure 2D), which record a period of basin desiccation (Legarreta, 1991; Gulisano and Gutiérrez Pleimling, 1995; Legarreta and Uliana, 1996).

Description: Unit 1 is 5.5–28 m thick and contains the Burro marker bed (Datum A) (Figure 2C), a light-grey indurated graded siltstone that is sharply overlain by light-grey fine-grained, planar-parallel laminated sandstone (Figure 4A). This unit is truncated by the basal surface of Unit 2 and is thinnest in the central sector of the exposure (see sections CML-9 to CML-16 Figure 3). Unit 1 comprises a heterolithic succession of planar-parallel laminated mudstones (F1) and thin-bedded (<0.1 m thick) normally-graded, well-sorted siltstones (F2) to very fine-grained sandstones (F3), and occasional medium-bedded structureless sandstones (F5) (Figures 4A,C). When traced from west to east, the thin-bedded sandstones show subtle lateral fining and thinning, transitioning from heterolithic succession to mudstone-prone succession. Unit 1 is rich in ammonites, belemnite rostrums and bivalves, as well as calcareous concretions (Damborenea, 1990; Gulisano and Gutiérrez Pleimling, 1995; Riccardi et al., 2011).

The central and eastern sectors contain a discrete stratigraphic interval that exhibits deformed bedding (Figure 4B). This interval is thickest (at least 10 m; Figure 3) in the central sector, where, internally, it exhibits an array of imbricated decametre-scale east-verging thrusts (offset <2 m) and associated drag folds. The thrusts originate from a bed parallel surface, leaving the underlying bedding undeformed (Figure 4B). In the eastern sector, a thin (∼5 m thick) interval of intense deformation is characterised by open folds and minor thrusts (offset <1 m) (Figure 4B). Unit 1 stratigraphy in the western sector lacks any deformation.

Interpretation: The laminated mudstones, graded siltstones and thin sandstone beds are interpreted as deposits of low-density turbidity currents (Allen, 1982; Trabucho-Alexandre et al., 2012; Könitzer et al., 2014), whereas the medium-bedded sandstones represent the deposits of medium-to high-density turbidity current (Talling et al., 2012). The laterally extensive character, mudstone dominance, and overall eastward (downdip) fining and thinning trend of thin-bedded sandstones of Unit 1 suggest deposition from low-energy sediment gravity flows in distal areas (e.g., Mutti, 1977), with possible distal lobe fringe deposits (Boulesteix et al., 2020). The discrete intervals of deformed bedding found in the upper parts of Unit 1 represent a post-depositional sheared zone linked to the overlying Unit 2.

Description: Unit 2 has an unconformable basal contact that truncates Unit 1 in the central sector (Figure 3). The relief of the basal contact is characterised by down- and up-stepping segments (ramps, >2°) linked by bedding-parallel segments (flats). The average thickness of Unit 2 is 20–30 m but can locally reach >47 m in the central sector and abruptly thins to <8 m towards the eastern and western sectors (Figure 3). This increase in thickness coincides with deeper erosion on the basal surface.

Unit 2 is characterised by a matrix-supported medium-grained muddy sandstone to sandy mudstone and is very poorly sorted throughout, ungraded, and with a chaotic distribution of outsized clasts (F14; Figure 5). Clasts range in character and size from granular quartz grains and rounded volcanic epiclasts to much larger megaclasts (>4.1–140 m long) (Hodgson et al., 2019) of either conglomeratic or heterolithic lithology (Figures 2C,D, 4F, 5). The chaotic distribution of polymictic clast encased into a muddy sand matrix is responsible for the block-in-matrix fabric (e.g. Ogata et al., 2012). Typically, conglomeratic megaclasts are rounded, elongated and weakly deformed (Figures 2D, 4F,G) and are clast-supported, with well-rounded to sub-angular clasts (0.03–1 m diameter) and fragments of thick-shelled bivalves (oysters; Figure 4D). These oyster-bearing conglomeratic megaclasts are preferentially located near the base of Unit 2 (Figures 2C,D, 3B). In contrast, heterolithic megaclasts are angular and characterised by internally folded packages of planar laminated and normally graded thin-bedded material (Figures 2C, 5) and preferentially distributed toward the top of Unit 2 (Figures 3, 5). This heterogeneity promotes a homogeneous matrix-rich texture in the middle division.

Apart from the irregular basal contact of Unit 2, thickness changes within the unit are strongly controlled by the rugose upper surface. Kilometre-scale wavelength (1–3 km) and metre-scale amplitudes (0.5–8 m) are responsible for a complex supra debrite topography.

Interpretation: The sedimentary characteristics of this unit, such as the chaotic distribution of (mega)clast floating onto a muddy sandstone to sandy mudstone matrix, suggest near-instantaneous deposition from a flow with high yield strength and buoyant support that could transport clasts up to 140 m long (Stow and Johansson, 2000; Mulder and Alexander, 2001). Using Datum A, the debrite formed a long-wavelength mounded top (Figure 3), attributed to the parental flow’s cohesive nature and en-masse freezing. We interpret Unit 2 as a cohesive debris flow deposit (Talling et al., 2012). The ramp and flat geometry at the base of Unit 2 indicate the debris flow’s erosive nature (e.g., Lucente and Pini, 2003). The similarity in composition between the heterolithic megaclasts and underlying Unit 1 suggests entrainment of deep-marine substrate blocks due to the shear stress exerted by the overriding debris flow (Van der Merwe et al., 2009; Watt et al., 2012; Ogata et al., 2014a; Hodgson et al., 2019). In contrast, the sand content in the matrix and the oyster-bearing conglomerate megaclasts suggest a shallow marine origin of the mass failure (e.g., Ogata et al., 2012). Alternatively, the megaclast bearing shells could come from the remobilization of older slope strata, including shallow-marine deposits (La Primavera Fm; Suárez and de la Cruz, 1997; Llambías and Leanza, 2005). The two distinct megaclast sources suggest long-distance transport of clasts and flow bulking through local substrate entrainment (e.g., Sobiesiak et al., 2016).

Description: Unit 3 (0–4 m thick) is composed of thick sandstone beds (0.5–2 m) with sharp, irregular concave-up bases and abrupt pinchout terminations, which result in a disconnected distribution of packages of wedge-shaped sandstone bodies (Figure 6) (see architecture section). Unit 3 is only present where Unit 2 is relatively thin (in the eastern-central and western sectors) and is absent in the central region where Unit 2 shows its maximum thickness (Figure 3). Where Unit 3 is absent, Unit 4 overlies Unit 2 (Figure 3). Unit 3 comprises bed types characterised by two main amalgamated divisions (lower and upper divisions) with some grain size breaks lacking any mudstone- and siltstone-rich bounding intervals (Figures 6A, 7). The basal interface of these sandstone bodies shows centimetre-scale undulations characterised by abundant load casts, semi-detached ball structures, and mudstone intrusions (diapirs and injectites) originating from Unit 2 (Figure 6G).

Two different types of thick-bedded amalgamated sandstone facies dominate the lower division, which varies along the transect. Grain-size breaks define amalgamation surfaces within sandstones. In the western sector, and more rarely in the eastern sector, the lower divisions are characterised by thick-bedded (0.5–2 m thick), structureless, weakly normally-graded, moderately-to poorly-sorted sandstones (F12). At bed bases, these sandstones comprise well-rounded (0.1–1 m diameter) mudstone clasts of low-sphericity and diffuse boundaries (mudstone clast type A), which show a coarse tail grading (Figures 6A,G). Locally, in the eastern sector, lower divisions of these sandstone bodies comprise thick-bedded, structureless, very poorly-sorted, more argillaceous sandstones with abundant mudstone clasts (0.1–1 m diameter) with very diffuse boundaries (mudstone clast type A), which are ungraded and randomly orientated (F13) throughout the encasing matrix (Figure 5F).

The lower division of Unit 3 sandstone bodies are overlain by an upper division (up to 2 m thick), which comprises coarse to very fine-grained, normally graded, moderately-to poorly-sorted sandstones (0.5–1.7 m) (Figure 6A). Banding can be developed throughout the bed or overlying a structureless division (Figure 6A). The banding is characterised by an alternation between lighter matrix-poor bands and darker matrix-rich bands that comprise bedding parallel millimetric mudstone clast with sharp boundaries (mudstone clast type B) (F8; Figure 6C). Contacts between bands are diffuse (Figures 6B,F). The spacing between the individual bands (0.5–2 cm) increases from the margin to central parts of the sandstone body (Figure 8), commonly showing rotation (Figure 6E). These sandstones develop symmetrical and asymmetrical convolute lamination at bed tops (predominant vergence towards NE; Figure 6D). Decimetre-scale long and centimetre-scale thick mudstone injections can be observed within this division (Figure 6B).

Interpretation: The wedge-shaped and deformed concave-up basal contacts of the sandstone bodies beds in Unit 3 are interpreted to reflect the interaction with the rugose upper surface and syn-sedimentary foundering of sand into the underlying mud-rich debrite (Figures 6, 9). Foundering is driven by instability due to the density contrast between the sand deposited above a less dense debrite (density loading) and lateral changes in sediment load (uneven loading) (Owen, 1987, 2003) produced by the short-wavelength rugosity of the upper surface.

The lack of sedimentary structures in the lower divisions of bed types recognised in Unit 3 is interpreted as a product of hindered settling from highly-concentrated gravity-flows, resulting in turbulence damping and rapid deposition (Talling et al., 2012), inhibiting any period of traction (Sumner et al., 2008). The normally-graded lower divisions were produced by incremental layer-by-layer deposition from high concentration gravity flows, such as high-density turbidity currents (sensu Lowe, 1982). In contrast, the thick-bedded argillaceous sandstones with ungraded mudstone clasts observed in the distal areas (eastern sector) are interpreted as moderate-strength cohesive debrites (sensu Talling et al., 2012). The decimetre-scale mudstone clasts (type A) were transported due to the matrix strength of the debris flows and their positive buoyancy with respect to the encasing matrix.

The lateral facies transition from high-density turbidites to moderate strength cohesive debrite suggest a flow transformation due to the entrainment of cohesive material from the underlying debrite (e.g., Kane and Pontén, 2012; Baker et al., 2017). The unconsolidated state of the debrite might have enhanced the substrate entrainment of decimetre-scale mudstone clasts (type A) and disaggregation (as indicated by the diffuse boundaries: Figure 6G), increasing the amount of mud and, therefore, the cohesiveness of the flow. Based on facies juxtaposition, the foundered sandstones can be subdivided into two different facies associations: 1) proximal and 2) distal, foundered sandstones facies associations (Figure 7). Both high-density turbidites and moderate strength cohesive debrites are characterised by rapid deposition (incremental deposition and en masse freezing, respectively), triggering the liquefaction of the fluid-saturated and unconsolidated upper surface of the debrite and foundering of sand (Figure 9). The undulations of the concave-up basal interface reflect complex interactions with the substrate: as the denser sand sank into the fluid-saturated muddy substrate, the buoyancy of mud promoted the syn-to post-depositional intrusion (mud diapirs and injectites) of the substrate into sandstones. The most advanced stage of foundering is observed when detached sand-balls develop (Figure 5; Owen, 2003; Tinterri et al., 2016).

In contrast, the banded sandstone characteristic of the upper divisions is interpreted to be formed under episodic near-bed turbulence damping at high rates of deposition (Lowe and Guy, 2000). The juxtaposition of the banded sandstones over the mudstone-clast bearing sandstones of the lower divisions suggests highly stratified flows, mixing and upwards transfer of centimetre-scale mudstone clast (type B) and the cohesive material from the disaggregation of the entrained decimeter-scale mudstone clast (type A). This enrichment in cohesive clayey material triggered the periodic suppression of turbulence and, therefore, banding development. The banding passes into convolute laminations towards the top, indicating moderate rates of deposition. The vergence of convoluted laminations suggests a syn-sedimentary shear-stress exerted by the overriding flow (McClelland et al., 2011; Butler et al., 2016) and flow-rebound produced by the underlying debrite relief (e.g., Tinterri et al., 2016).

Description: Unit 4 (10–27.3 m thick) has a sharp and concordant contact with the underlying Unit 2 and Unit 3 (Figure 3). It comprises two subunits: a lower heterolithic interval (Unit 4A) and an upper sandstone-prone interval (Unit 4B; Figure 2C).

Unit 4A is thin-to medium-bedded (0.01–0.5 m; Figure 4H) heterolithic succession (F1, F2, F3 and F5) with a maximum thickness of 22 m, thinning to 8 m in the central sector above where Unit 2 is thickest (Figure 3). Most of the thin-beds (0.01–0.1 m thick) are fine-to medium-grained normally-graded sandstones, matrix-poor, moderately well-sorted, and structureless with common planar-parallel lamination and/or starved-ripple lamination near bed tops (F3; Figure 4C). Palaeocurrent measurement shows a consistent flow trend towards the NE (Figure 9). Unit 4A also comprises coarse-to granular normal-graded sandstones, relatively low matrix content with common grain size-breaks (0.07–0.2 m thick), erosive bases and sharp-planar tops (F4; Figure 4D) and two medium-bedded matrix-supported conglomerates with sandstone clasts (F14; 0.25 and 0.35 m thick, respectively) that pinch out towards the central sector (Figure 8). One of these thin gravelly beds, which lacks any lateral thinning or fining trend (Figure 4I), was traced across the exposure (gravelly marker bed; dashed red line in Figure 8). In addition, a 0.15 m thick tuff layer (Figure 4J) was also used for correlation purposes (tuff marker bed; dashed white line in Figure 3). The medium-bedded sandstones (0.1–0.5 m thick) are structureless, ungraded, with planar-parallel and convolute lamination at bed tops, except one that shows cross-bedding (F6; Figure 4D). These sandstones have sharp bed bases and tops and lack mudstone clasts. In the eastern sector, Unit 4A is dominated by thin-to medium-bedded heterolithic succession that lacks any gravelly (F5) or matrix-supported conglomerate beds (F14).

Unit 4B (5.7 m thick in the western sector) thins eastwards along a 4.3 km transect until it pinches out, where Unit 2 is thickest (Figure 3). In the western sector, it dominantly comprises medium- (F5; Figure 4K) to thick-bedded sandstones (F11; Figure 4K), with less common “bipartite” sandstone beds (F9 and F10) composed of a matrix-poor lower division and a matrix-rich upper division with mudstone clasts. Conformable bases and sharp tops characterise the thick-bedded sandstones (0.5–1.2 m). Where the thick-bedded sandstones are not amalgamated and are intercalated centimetre-thick beds of fine-grained material (F1 and F2), bed bases are loaded locally. The thick-bedded sandstones are normally graded from medium to fine sand, well-sorted with rare centimetric mudstone clasts at the bed top. Soft-sediment deformation structures, such as centimetre-scale flames with NE vergence, are also common at bed bases and along amalgamation surfaces (Figure 4L). Banded sandstones are medium-bedded (0.1–0.5 m), fine-to medium-grained, and characterised by alternating between light- and dark-coloured bands, ranging from 0.2 to 2 cm thick (F7; Figure 4L). Both band types show a similar maximum grain size, although the darker bands are matrix-rich, and light bands are matrix-poor. Banding is generally sub-parallel to bedding. Although banded sandstones are more commonly associated with thick-bedded structureless sandstones, the banded sandstones can be individual event beds, with banding above the structureless basal division. The medium-bedded bipartite sandstone beds (0.1–0.5 m) consist of a medium-grained, matrix-poor and structureless lower division, which is overlain by a fine-grained matrix-rich upper division characterised by poor sorting and abundant mudstone clasts (0.05–0.3 m) with low sphericity and variable roundness (F9 and F10; Figure 4M). The lower and upper divisions show a gradual upwards increase in matrix content rather than across a sharp boundary. When Unit 4B is traced eastwards towards the central sector, the sandstone package transitions into a few thin-beds (0.1 m thick) of weakly graded, very poorly-sorted matrix-rich sandstone, lacking the mudstone clasts observed in western areas. Unit 4B is absent in the eastern sector.

Interpretation: In Unit 4A, the thin sandstone beds showing planar and cross ripple laminations support an interpretation as low-to medium-density turbidites (Talling et al., 2012). The starved-ripple lamination observed in thin-bedded sandstones is interpreted as the reworking of sand deposited by dilute flows with low sedimentation rates (Talling et al., 2007; Jobe et al., 2012). The intercalation of thin-bedded sandstones with finer-grained deposits suggests a lobe fringe environment (Lobe fringe facies association, Figure 5) (Prélat et al., 2009; Spychala et al., 2017b). The abundant coarse-grained to gravelly thin-bedded sandstones in the western sector record intermittent energetic coarse-grained flows, suggesting sporadic sediment bypass processes (Stevenson et al., 2015). However, the low matrix content within the granular beds suggests a sediment source area where only coarse-to granular grain size was available. The intercalation of such different facies suggests the juxtaposition of depositional environments of contrasting energy and/or different sediment sources. Either scenario could be possible given the complex sediment routing patterns and multiple transverse or axial sources available in the Neuquén Basin during the early post-rift setting (Vicente, 2005; Privat et al., 2021) and by analogy to other post-rift settings (e.g., Lien, 2005; Fugelli and Olsen, 2007; Hansen et al., 2021). The mass failure would trigger a new coarse-grained source due to slide scar position and geometry, promoting intermittent sand supply to an otherwise sand-starved environment (see “Origin and role of the Mass-wasting process as a trigger for turbidite systems development” in discussion). The downdip variability in the thickness of Unit 4A (from 22 to 8.5 m thick), reduction in gravelly sandstone content and the stratigraphic thinning between the granular marker bed (red dashed line in correlation) and the top debrite (Unit 2) reveals the existence of subtle relief on the debrite surface (Figure 3B). Furthermore, the two poorly sorted ungraded muddy sandstones, which are interpreted to be debrites due to their chaotic distribution of clast within the argillaceous matrix, also pinch out towards the central sector. The ripples and convolute laminae with SW vergence (Figure 3) contrast with the consistent NE paleoflow, suggesting local flow deflection (cf. Tinterri et al., 2016) in the central sector, where the debrite relief is highest, indicating the interaction between sediment gravity-flows and the upper surface of the debrite.

Massive medium-to thick-bedded deposits of Unit 4B are interpreted as high-density turbidites formed by incremental layer-by-layer deposition with high aggradation rates (Kneller and Branney, 1995), interpreted to represent proximal lobe axis environments (lobe axis facies association; Figure 5) (e.g., Prélat et al., 2009; Kane et al., 2017). The location of these facies in the westernmost sector, and the palaeoflow measurements, suggests that the western sector was relatively proximal. Banded sandstones represent the deposits of mud-rich transitional flows formed by tractional reworking (Stevenson et al., 2020). The bipartite beds consisting of a basal structureless to planar laminated sandstone division, overlain by a linked mudstone clast-rich upper division are interpreted as hybrid event beds (HEBs), formed from transitional flows deposited under high-deceleration rates (Haughton et al., 2009; Hodgson, 2009; Kane and Pontén, 2012) in more distal environments than the banded sandstones (Stevenson et al., 2020). The gradual and diffuse boundary between the basal turbidite and the upper debrite suggest vertical segregation of particles within the cohesive flow (Kane et al., 2017). The facies evolution of Unit 4B from proximal (western sector) to distal (eastern sector) of thick-bedded sandstones into hybrid event beds likely represents the downdip transition from lobe-axis/off-axis environments (lobe axis facies association: Figure 5) (sensu Prélat et al., 2009) into lobe-fringe environments (lobe fringe facies association; Figure 5) (e.g., Spychala et al., 2017a; Kane et al., 2017), persisting until the frontal/oblique pinchout (e.g. Hansen et al., 2019).

Using Datum A (“Burro” marker bed), the upper surface of the >9.6 km long debrite forms a broad convex-up relief that reaches a maximum in the central section coincident with the deepest incision (at least 22.5 m of erosional relief; Figure 3B). The spatial association of the thickest part of the debrite with the deepest incision support a genetic link between the geometry of the flat-ramp-flat shaped basal shear zone and the mounded top. The morphology of the basal surface can buttress material translated downslope and develop positive topographic features, such as pressure ridges (Moscardelli et al., 2006; Bull et al., 2009). Bed-by-bed correlation within Unit 4A shows that where the upper surface of the debrite develops the highest relief (∼8 m of positive relief with respect to the western sector), Unit 3 is absent, and it is overlain by Unit 4A (Figure 3), showing a laterally continuous stratigraphic interval with metre-scale thickness variations (Figure 8). Unit 4A thins from 22 m (CML-1) and 13 m (CML-2) to 6 m (CML-12) across the highest part of the debrite (Figure 3B). The lower part of Unit 4A pinches out in the central sector, developing onlaps of individual beds, and supporting the existence of a gentle relief (Bakke et al., 2013; Soutter et al., 2019). In contrast, the upper part of Unit 4A shows tabular architecture with a lateral continuity of over 7 km.

Unit 4A is overlain by Unit 4B, which shows a progressive thinning of the submarine lobe from the western to the central sector over 5.6 km, from 5.7 m (CML-3) to 1.7 m (CML-12) and 1 m (CML-14) with a mean thinning rate of 0.9 m/km. The submarine lobe pinches out between CML-14 and CML-22 (<2 km), interfingered with unit 4A (Figure 3B). The lack of onlap geometries against underlying deposits and subtle thinning rates consistent with unconfined settings (e.g., Prélat et al., 2009) suggests a lack of a pronounced pre-existing relief. However, the coincidence of lobe pinch-out in the area where the debrite relief is highest and where the underlying Unit 4A is thinnest might reflect subtle residual relief.

The steeply-dipping unconformable base, internal deformation and abrupt thickness changes of Unit 3 sandstones contrast with their flat and conformable tops (see stereoplots in Figure 6). These sandstone bodies can be subdivided into three different types by their architecture.

As submarine landslides travel across the seafloor, they exert shear stress on the substrate, coupled with significant over-pressure (Bull et al., 2009; Hodgson et al., 2019; Payros and Pujalte, 2019). This leads to substrate entrainment (Eggenhuisen et al., 2011; Hodgson et al., 2019) and/or deformation (Butler and McCaffrey, 2010; Watt et al., 2012; Dakin et al., 2013; Ogata et al., 2014b). The debris flow (Unit 2) incised at least 22.5 m into the substrate (Unit 1; Figure 3). In the central sector, the basal shear surface forms ramps (up to 800 m long, >2°) and flats (up to 1,550 km long; Figure 3B; between the logs CML-9 and CML-10) (see Lucente and Pini, 2003; Frey Martinez et al., 2005 for flat-ramp-flat geometry). The stress applied to the substrate during the emplacement is accommodated by both stratigraphic intervals consisting of deformed packages (basal shear-zone) and interfaces consisting on a plane (basal shear-surface) (Alves and Lourenço, 2010), such as the discrete basal shear zone located in upper Unit 1. The absence of contractional features in the deposits underlying Unit 1 supports the deformation as the result of the shear stress produced by the emplacement of the debrite rather than tectonism. The basal shear zone has variable thickness and deformation styles. It is absent in the western sector, whereas erosion and deformed intervals record a high degree of basal shear stress in the central sector (Figure 3B). In the central sector, the deformed package (up to 10 m thick) is characterised by decametre-scale thrusts with metre-scale offsets and drag folding (Figure 4B). The predominance of imbricate thrusting over folding, and lack of internal disaggregation within the package, indicate competent substrate rheology (e.g., Van Der Merwe et al., 2011). The eastward vergence of the compressional structures (Figure 4B) indicates an eastward emplacement direction for the debris flow (Twiss and Moores, 1992), consistent with the palaeoflow indicators in the bounding strata.

The thrusting is attributed to bulldozing by the entrenched debris flow (e.g., Jackson, 2011; Hodgson et al., 2019; Payros and Pujalte, 2019), representing the initial stage of substrate entrainment. Entrainment of megaclasts into a debris flow has been reported in other systems in the subsurface (Moscardelli et al., 2006; Alves and Cartwright, 2009; Sawyer et al., 2009; Dakin et al., 2013; Ortiz-Karpf et al., 2015; Soutter et al., 2018; Nwoko et al., 2020a) and more rarely at outcrop (Ogata et al., 2014b; Sobiesiak et al., 2016; Hodgson et al., 2019; Cumberpatch et al., 2021). The progressive increase in thickness and degree of strain along the basal shear zone of Unit 2 and the enrichment in rafted heterolithic megaclasts (Figures 3B, 5) suggest downdip evolution of the debris flow, which might have affected the parental debris flow rheology (e.g., Hodgson et al., 2019; Payros and Pujalte, 2019) and bulking of the flow (Gee et al., 2006; Alves and Cartwright, 2009; Butler and McCaffrey, 2010; Hodgson et al., 2019). The preferential location of heterolithic megaclast towards the top of Unit 2 might be related to internal granular convection cells created along with the debris flow, enhanced by the buoyancy of less dense rafted megaclast compared to debrite matrix (Hodgson et al., 2019) and kinetic sieving (Legros, 2002). In contrast, conglomerate megaclasts are always found at the base of the debrite due to their higher density than the surrounding debrite matrix.

The absence of Unit 3 sandstones over the thickest part of the debrite suggests that the sediment gravity flows were strongly stratified and influenced by the debrite relief (Figure 8). The sandstone bodies are also disconnected at finer scales, revealing short wavelength (metre-scale) and amplitude (decimetre-scale) rugosity on the debrite surface. The existence of simultaneous short wavelength and amplitude rugosity superimposed on a large-scale wavelength relief on the upper surface of an MTD has also been reported by Armitage et al. (2009), defined as “surface-topography hierarchy,” in the Cretaceous Tres Pasos Formation at the Sierra Contreras (Chile) and by Fairweather (2014) in Carboniferous Paganzo Basin at Cerro Bola (Argentina). In this study, the deposition of sand in pre-existing lows filled the short-wavelength rugosity and triggered the loading of individual sandstone bodies onto the mud-rich debrite (See “Small-scale architecture: Foundered sandstones” sections), leaving the large-scale relief underfilled (Figures 3, 8). The foundering process is evidence of substrate liquefaction and highlights the dynamic interface between the debrite and subsequent flows and their deposits. A similar scenario was proposed by Van der Merwe et al. (2009), Van der Merwe et al. (2011) in the Vischkuil Formation in the Laingsburg depocentre (Karoo Basin).

The ability of supra MTD rugosity to pond turbidity currents travelling across their upper surface is a well-known phenomenon (Kneller et al., 2016). However, the presence of Unit 3 foundered sandstones up-dip and down-dip of the debrite high (Figure 3B), and its consistent NE paleocurrent trend, suggest connected sediment transport routes across the debrite with no evidence of flow ponding or stripping (e.g. Armitage et al., 2009; Fairweather, 2014). The highly-stratified grown-hugging parental flows of Unit 3 would have been ponded in proximal parts (western sector) if a fully enclosing topography existed given their reduced ability to surmount obstacles (Al-Ja'Aidi et al., 2004; Bakke et al., 2013), resulting in sand starvation over the debrite in distal settings (Sinclair and Tomasso, 2002; Kneller et al., 2016). The overlying Unit 4A can be traced laterally across the study area, with metre-scale thinning where the debrite is thickest (see CML-12; Figure 3). Apart from this, the advancing onlap geometries of the thin beds and the divergence in the overall NE-orientated paleocurrents (rose diagram Unit 4; Figure 3) indicate the progressive healing of the large-scale wavelength debrite relief, with some deflection of turbidity currents (Figure 10).

The thin sandstone beds of the upper part of Unit 4A healed the debrite high. However, the gravelly beds thin and fine from proximal to distal areas (western to eastern sectors), and the two debrites pinch out in proximal areas (western sector), suggesting subtle remnant topography (Figure 8). The different lateral continuity of individual beds is explained by different rheologies of individual sediment gravity flows, which affect the flow efficiency (Al Ja’Aidi et al., 2004). Cohesive debris flows are more influenced by irregular relief, while low-density turbidity currents are less affected by seafloor topography (Bakke et al., 2013; Soutter et al., 2019). This suggests that laterally continuous thick accumulations of lobe fringe deposits can develop on gentle topographies, while the submarine lobes’ axial parts were restricted to lower relief areas. The interaction of thin-bedded turbidites successions with gentle topography has also been reported in other deep-water settings (i.e., “aggradational lobe fringes”; Spychala et al., 2017b). The deposition of lobe fringe successions reduced confinement, which enabled the deposition of the Unit 4B submarine lobe.

The submarine lobe is characterised by a progressive thinning and fining, developing pinch-out geometries and interfingering with Unit 4A in the area where the relief of the debrite is highest. The development of pinch-out geometries over the areas where the debrite shows a mounded relief and where the Unit 4A onlaps and thins suggests that the relief was not wholly healed with the deposition of Unit 4A and affected the parental flows of Unit 4B.

One explanation is that the exposure exhumes the Unit 4B lobe obliquely, with the medium-to thick-bedded lobe axis deposits in the westernmost sector (CML-1 to CML-4) transitioning into an HEB-dominated fringe, being highly impacted by gentle seafloor topography (Soutter et al., 2019; Privat et al., 2021). Alternatively, the seafloor relief could have promoted the modification of flow pathways and deflection of flows, thus changing the downdip orientation (Figure 8). All these scenarios suggest a confined and uncontained (see Southern et al., 2015) lobe-type depositional system. The precise dispersal pattern of the flows remains unknown due to the outcrop limitations. Nonetheless, the documented stratigraphic evolution reveals that long-lived debrite relief and progressive healing by deposition of aggradational lobe fringes enabled the progradation of sand-rich submarine lobe, albeit with changes in flow rheology the bed style and element-scale pinchout (Figure 8).

Another explanation is that the debrite relief in the central sector might have been rejuvenated through volume changes in the debrite due to differential compaction by loading the lobe itself in the proximal sector. However, it seems unlikely that the deposition of a 5.7 m thick lobe could promote a volume loss in an 8–47 m thick debrite, given that both units are separated by 8–22 m thick thin-bedded interval. In contrast, another two hypotheses are deemed more plausible; fluid loss and fault-controlled mechanical subsidence. Fluid loss-controlled evacuation could have promoted differential subsidence of the upper surface of the debrite and overlying units (e.g., Browne et al., 2020). Alternatively, given the early post-rift setting, mechanical subsidence by an east-facing and N-S striking fault (Manceda and Figueroa, 1995) could have generated more accommodation in the western part of the study area (see Cristallini et al., 2006). However, this implies a very localised and rapid reactivation, and there is no other evidence for post-rift tectonism identified.

The emplacement of the >9.6 km long and erosional debrite in the Chacay Melehue depocentre reflects an abrupt change in sedimentation patterns, which were previously dominated by dilute mud-rich flows (Unit 1). The first significant sand influxes in the depocentre for ∼6 Myr since the Aalenian (interval II of Gulisano and Gutiérrez Pleimling, 1995) are recorded by the sand-rich deposits (Unit 3) immediately overlying the debrite (Figure 10). The juxtaposition of sand-rich turbidites over debrites (metres to hundreds of meters thick) have been reported in other systems (Kleverlaan, 1987; Labaume et al., 1987; Payros et al., 1999; Fallgatter et al., 2017). These authors suggest that the debris flow underwent a period of mixing with ambient water, leading to the generation of an overriding co-genetic turbidity current. The foundering phenomenon reported here reveals a close spatiotemporal relationship between the debrite emplacement (Unit 2) and overlying sandstone deposition (Unit 3). An alternative mechanism is that the mass-failure event altered the basin margin physiography such that a sand source was captured. Mass-wasting processes responsible for the evacuation of material from shelf edge and upper slope areas alter the bathymetric configuration of basin margins and promote the funnelling of sediment stored in shallow marine environments through slide scars (e.g., Moscardelli and Wood, 2008; Kneller et al., 2016; Steventon et al., 2020) (Figure 11). The role played as a trigger mechanism for sand delivery into deep-water setting by the mass-wasting event is a plausible scenario given the sand-starvation recorded in the coeval deposits of Los Molles Fm along the eastern margin of the Chacay Melehue depocentre (Veiga et al., 2013).

Given the palaeoflow and kinematic indicators, the thickness patterns of the studied units (Figure 3B), and previous studies on sediment supply from the volcanic arc (Vicente, 2005), we propose that the mass failure originated to the west of Chacay Melehue, where a major syn-rift fault is located close to the volcanic arc (<30 km; Manceda and Figueroa, 1995; De La Cruz and Suarez, 1997; Vicente, 2005). The role of the western volcanic arc as a source area for the early post-rift sediment supply in the Chacay Melehue depocentre is supported by the southeastwards directed paleocurrents measured in the Aalenian turbidite system at the base of the Los Molles Formation (Vicente, 2005; Figure 2) and the abundance of pyroclastic deposits within Los Molles Formation stratigraphy (Zöllner and Amos, 1973; Rosenfeld and Volldaeimer, 1980; Gulisano and Gutiérrez Pleimling, 1995; De La Cruz and Suarez, 1997; Llambías and Leanza, 2005). The oyster-bearing conglomerate megaclast and well-rounded volcanic epiclast within the matrix of the debrite reflect long-lived reworking in shallow-marine settings prior to the mass failure, suggesting a shallow-water origin or remobilization of older slope strata, including shallow-marine deposits. This could represent the downslope transfer of sand following the collapse of reworked volcaniclastic deposits along the magmatic arc (Figure 9). The evolution from the initial mass-wasting sediment supply responsible for erratically distributed foundered sandstone bodies (Unit 3) to a more mature system with the subtle distribution and diversity of lobe architectural elements (Unit 4) reflects the evolution to a more organised sediment supply system (Figure 10). This was abruptly superseded by a return to sand-starved conditions dominated by dilute mud-rich flows and hemipelagic deposition until the end of the Lower Callovian (Gulisano and Gutiérrez Pleimling, 1995).