Evolution from syn‐rift carbonates to early post‐rift deep‐marine intraslope lobes: The role of rift basin physiography on sedimentation patterns

The stratigraphic architecture of Early Jurassic strata exposed along a >10 km long transect in the Chachil Graben, an exhumed marine rift depocentre in the Neuquén Basin (Argentina), provides insights into the sedimentological and stratigraphic expression of the syn‐rift to post‐rift transition. A change from syn‐rift intrabasinal carbonate to post‐rift extrabasinal siliciclastic sedimentation is recorded, as well as variations in sediment supply and dispersal patterns across rift‐related topography. The late syn‐rift was marked by a transgression and development of a shallow‐marine carbonate system, including carbonate platform deposits perched on fault‐block highs and periplatform deposits accumulated in fault‐block lows, which overlies continental volcano‐sedimentary syn‐rift deposits. Differential subsidence and basin deepening induced retrogradation of the carbonate system, which was progressively drowned and overlain by organic‐rich calcareous mudstone that draped across rift structures at the onset of the early post‐rift. The first extrabasinal siliciclastic influx led to progradation of an early post‐rift intraslope lobe complex into the graben, which is associated with kilometre‐scale clastic injectites. The depositional architecture, facies distribution and pinch‐out style of intraslope lobes record the effects of an inherited compaction hinge, which acted as an oblique counterslope to sediment gravity flows. The occurrence of combined‐flow bedforms, widespread erosion, and limited facies segregation across lobes bearing different hybrid event bed types, is in sharp contrast to sedimentological characteristics of existing intraslope lobe models. Documentation of the syn‐rift to post‐rift transition stratigraphy permitted identification of changes in thickness and facies resulting from the passive infill of inherited topography with early post‐rift differential compaction. This architecture contrasts markedly with those developed during syn‐rift normal faulting. Furthermore, the influence of local inherited topography on the development of early post‐rift lobes is key to improve subsurface prediction of sandstone distribution and quality during assessment of hydrocarbon reservoirs and carbon storage sites.


INTRODUCTION
The syn-rift to post-rift transition is associated with a change from active crustal stretching, typically accommodated by normal faulting and rapid subsidence, to slow thermally-induced subsidence associated with cooling of the crust and the cessation of normal faulting (McKenzie, 1978;Ziegler & Cloetingh, 2004). Depending on the magnitude and distribution of inherited rift topography, the location and rate of thermal subsidence, and eustatic and climatic changes, the rift may receive extrabasinal sediment almost immediately in the early post-rift phase, or this may be delayed until later in the post-rift (Ravn as & Steel, 1998;Lien, 2005;Soares et al., 2012;Yu et al., 2013;Jarsve et al., 2014;Henstra et al., 2016;Bal azs et al., 2017).
Detailed documentation of strata recording the syn to post-rift transition is needed to evaluate controls such as inherited rift topography, and changes in sediment source area and/or routing pathways, to improve existing conceptual models for rift basin-fills. However, the syn-rift to postrift transition is difficult to resolve in the subsurface, due to typically low-resolution of seismic reflection data and sparse well coverage (Kyrkjebø et al., 2004;Lien, 2005;Zachariah et al., 2009;L opez-Gamund ı & Barragan, 2012;Jarsve et al., 2014;Lohr & Underhill, 2015). Because of this, little is known about the detailed architecture, facies distribution and development of bed-scale heterogeneity, as well as the termination style of deep-marine early post-rift lobes against inherited rift topography, all of which have been documented in the subsurface (Argent et al., 2000;Martinsen et al., 2005;Milton-Worssell et al., 2006;Moscardelli et al., 2013;Dodd et al., 2019). Outcrop-based studies of ancient deep-water rift basin-fills can provide key information on sub-seismic stratigraphic architecture and facies distributions during the syn-rift to post-rift transition, but often lack detailed descriptions of early post-rift, sand-rich systems (Surlyk & Korstg ard, 2013;Yu et al., 2013;D'Elia et al., 2015;Hadlari et al., 2016).
This study investigates a rare example of an exhumed marine rift depocentre, the Chachil Graben, south-western Neuqu en Basin, west-central Argentina. The aim is to document the transition from syn-rift carbonate to post-rift siliciclastic sedimentation recorded in the Chachil depocentre during the Early Jurassic. The main objectives are to: (i) analyze the tectono-stratigraphic architecture of the depocentre-fill; (ii) provide a detailed characterization of early post-rift intraslope lobes; and (iii) discuss the impact of local and regional controls on sedimentation in order to refine the interpretation of the syn-rift to post-rift transition in the Chachil Graben and improve rift basin models. A model is presented showing the organization of sedimentary systems across the evolving rift topography, which incorporates changes in sedimentation regime, and supply and dispersal (i.e. axial, or transverse, to structural strike) during the syn-rift to post-rift transition. The influence of syn-depositional relief on the architecture, facies distribution and pinch-out style of early post-rift intraslope lobes can be used to improve hydrocarbon reservoir characterization, which can be otherwise challenging in the subsurface.

GEOLOGICAL SETTING AND STRATIGRAPHY
The Neuqu en Basin, Argentina (36°S to 40°S) formed along the south-western convergent margin of the Gondwana-South American plate, being bound by the Andean volcanic arc to the west, the Sierra Pintada belt to the north-east, and the North Patagonian Massif to the southeast (Fig. 1A). The Neuqu en Basin underwent intracontinental rifting from Late Triassic to Early Jurassic, followed by subduction of the proto-Pacific plate and thermal subsidence from Early Jurassic to Early Cretaceous; foreland basin evolution with compression and inversion characterized the Late Cretaceous -Cenozoic period (Vergani et al., 1995;Legarreta & Uliana, 1996;Franzese & Spalletti, 2001;Howell et al., 2005).
Transgression of volcanic rift depocentres occurred during the Early Pliensbachian along the southern margin of the Neuqu en Basin, with flooding from the Panthalassic Ocean and formation of an epeiric sea along the south-western Gondwana margin (Damborenea et al., 2013;Leanza et al., 2013). The Early Pliensbachian also recorded the onset of subduction with trench rollback, which controlled extension and back-arc subsidence of the Neuqu en Basin bounded to the west by the Early Andean volcanic island arc (Franzese & Spalletti, 2001;Mpodozis & Ramos, 2008). Initial development of Early Jurassic marine depocentres was strongly influenced by the rift topography and diachronous syn-rift to post-rift transition until the Middle Jurassic (Aalenian) (Legarreta & Uliana, 1996;Burgess et al., 2000;G omez Omil et al., 2002;Veiga et al., 2013). Since Middle Jurassic marine depocentres merged together into a single broad back-arc depocentre controlled by post-rift thermal subsidence (Gulisano & Guti errez-Pleimling, 1995;Vergani et al., 1995;Legarreta & Uliana, 1996;Franzese & Spalletti, 2001). The Early to Middle Jurassic evolution of the Neuqu en Basin is recorded by the Cuyo Group (Gulisano et al., 1984). It comprises two second-order depositional sequences separated by the Toarcian -Aalenian boundary: the Early Jurassic Lower Cuyo Group and Middle Jurassic Upper Cuyo Group (Gulisano & Guti errez Pleimling, 1995;Vergani et al., 1995;Legarreta & Uliana, 1996;Burgess et al., 2000). This paper focuses on the Early Jurassic transgressive deposits of the Lower Cuyo Group, represented in the study area by the Chachil Formation (Weaver, 1942) and the Lower Los Molles Formation (Weaver, 1931). The aim is to document the local syn-rift to post-rift transition record in the Chachil Graben and the effects of inherited early post-rift relief on the characteristics of intraslope lobes.

Chachil Graben
The study area corresponds to the Chachil Graben located in the south-western Neuqu en Basin, which forms the exhumed western part of the Huincul High (Fig. 1A). The Huincul High is an ENE-WSW oriented intraplate structure characterized by a series of small NNW-SSE to NW-SE trending half-grabens and grabens inverted during the Andean compression (Vergani et al., 1995;Franzese et al., 2006;Garc ıa Morabito et al., 2011;Muravchik et al., 2011Muravchik et al., , 2014. The Huincul High formed the southern shelf-slope margin of the Neuqu en Basin during the Early Jurassic, where a series of marine rift depocentres (see Fig. 1A) developed with turbidite fan systems across complicated rift topography (e.g. G omez Omil et al., 2002;P angaro et al., 2009).
The Chachil Graben is a NNW-SSE trending depocentre about 10 km wide and at least 15 km long, although post-rift cover and Cenozoic volcanic rocks overlying syn-rift deposits hamper the definition of its eastern margin (Fig. 1B). Its northern margin was controlled by the Chihuido Bayo fault system, which strikes north-east/south-west and dips to the south-east. The Chihuido Bayo fault system extends for 15 km south-east from the Cerro Chachil, where it bounds the 5 km wide horst that forms the Southern Chachil Graben margin ( Fig. 1B and C). The Lapa Formation, which represents the syn-rift infill of the Precuyano Cycle in the study area (Fig. 2), is <400 m thick along the Southern Chachil Graben margin and thickens up to 2 km northward in the graben centre (Franzese et al., 2006). Thickness changes in the Lapa Formation were controlled by NNW-SSE and minor NE/NNE-SW/SSW striking faults, respectively parallel and oblique to the Chihuido Bayo fault system along the main Southern Chachil Graben margin (Fig. 1B). North-west/south-east and NNW-SSE striking, north-east or south-west dipping late syn-rift faults subdivide the immediate hangingwall of the Southern Chachil Graben margin into several intra-basin highs (Morro del Aguila, Puesto Alfaro, El Luchador and Paine Milla faultblocks) and intervening lows (Mirador de Chachil and Pic un Leuf u fault-blocks) ( Fig. 3A and B). The variability in the trend of rift faults results from the interaction of the main northeast/south-west directed extensional stress field with pre-rift Palaeozoic basement fabrics that Fig. 2. Synthetic stratigraphic column for the Chachil Graben, representing thickness of lithostratigraphic units, changes in sediment composition (see heading 'SC'; note that '+' indicates an increase and '-' indicates a decrease in non-dimensionalized sediment content) and biota, source contribution and interpretations of tectonic stages. Numerical Early Jurassic ages from Ogg et al. (2016); (NC) nannofossil chronozones (Ballent et al., 2011); Standard European Ammonite Biozone (EAB) and Andean Ammonite Biozone (AAB) numbers (Riccardi, 2008) (Franzese & Spalletti, 2001;Mpodozis & Ramos, 2008;Cristallini et al., 2009). In the study area, the later eastwest Andean compressional stress field was oblique to existing rift fault trends and only resulted in partial inversion of rift structures (P angaro et al., 2006). Herein it is observed that this is expressed in the Chachil Graben with development of small-scale east/north-east dipping folds and thrusts affecting the Lower Los Molles Formation and large-scale thrusts reactivating pre-rift basement structures.

Data and methodology
A >10 km long north-east/south-west striking exposure belt into the Chachil Graben permits the detailed analysis of sedimentology and stratigraphic architecture, with a focus on stratal geometries, facies and thickness changes of both carbonates, and intraslope lobe systems, across rift structures. The dataset comprises detailed geological field mapping (Fig. 3) (Fig. 4). Present-day mean bedding orientations were calculated for each unit from all measurements taken in a given zone, to constrain an average tectonic dip, and dip direction, for each unit in each zone ( Fig. 4)   rift-related structures and disturbance of the depositional architecture was limited (Fig. 5). Locally, in the Zone NE, the steeper dips measured are related to structures that formed during later inversion tectonics across the El Luchador fault-block (Figs 3 and 4). Note that log sections 11 and 12 were not included in this analysis of structural dip given the paucity of data compared to other sections. Palaeocurrent measures (Fig. 5) were primarily collected from: sole marks including grooves and flute casts, ripples, and dune-scale cross-bedding, and were plotted in rose diagrams to reconstruct the palaeoflow pathways with bedding restored using stereonet software. Maps and cross-sections ( Fig. 3) are based on the present-day structural and topographic configuration (i.e. no back-stripping or structural restoration). The lack of precise data including palaeobathymetric or palaeoecological proxies, volumes of eroded sediment, local eustatic curve, and flexural rigidity of the lithosphere, do not allow back-stripping to be undertaken. Furthermore, the nature of the outcrops that limit the collection of three-dimensional data, the absence of a consistent horizontal datum in all vertical sections, and the changing stress field orientation through time (P angaro et al., 2006), hinders construction of balanced cross-sections and palinspastic restoration (cf. Spikings et al., 2015). Although absolute (original) depositional dips and dip azimuths cannot be determined, the use of the depositional architecture, bedding dip orientation changes, unconformities, thicknesses and palaeocurrents across rift fault-blocks, permitted the interpretation of synsedimentary variation in accommodation and sediment dispersal across faults, with healing of rift topography. Marker beds used for stratigraphic correlation were identified in Unit 4 ( Fig. 5) and primarily consisted of extensive sandstone packages that could be walked out for several kilometres; these were deposited across a less complicated seabed topography compared to other units. The ca 5 km exhumed belt of sandstone packages corresponding to the intraslope lobe complex represents the onlap margin of a broader sandy depocentre that continues in the subsurface to the south-east of the Chachil Graben. Physical correlation of sandstone packages and injectites between logs was also constrained with Uncrewed Aerial Vehicle (UAV) photographic panels.

Description
Facies Association 1 (FA1) consists of well-stratified, thin to medium-bedded silicified carbonate  Fig. 3A), showing relationships between units in the Chachil Graben, along with average discordance angles corresponding to the location of the logs (see Fig. 4) and palaeocurrent measurements. Each rose diagram shows the detail of sole marks (in red), and ripples and dune-scale cross-bedding current directions (in green). Marker beds used for correlation are shown with dashed black lines, and correspond to depositional sandstones and debrites (for specifics see the detailed architecture of intraslope lobes in Fig. 9), and the top of the clastic sill body (see Fig. 10). Note that there is uncertainty in lobe correlation between logs 1 and 2, and logs 5 to 9, due to basalt cover (see map Fig. 3A). The detailed architecture of the Lapa Formation is not represented and vertical offsets on faults are approximate. Orientation of sections is shown on the map in Fig. 3A.  (Damborenea & Manceñido, 1979Armella et al., 2016).

F2c
Calcarenite Grainstone, massive, dominated by well-rounded Ms to Fs sized allochems and quartz grains and thin shells (bivalves).

F4c
Calcareous bioclastic sandstone MS normally graded Ms to Fs with low angle planar and current-ripple laminations and with abundant broken skeletal material.

F5.1c
Graded sandy siltstone Normally graded sandy Si with sandy low-angle laminations and starved current-ripples.

F5.1d
Structured finegrained sandstone MS crudely to normally graded Ms to VFs with planar or undulose laminations and current-ripples, locally carbonaceous-rich.

F5.2a
Massive mudpoor sandstone PS Cs-Ms with clay-poor matrix, massive or with subtle normal grading and bearing few deformed subrounded to subangular Mds clasts (0.5-8.0 cm) near base or throughout. Mud-rich bed-waves (decimetre-scale) with low-angle laminae (5-10 cm high and wavelength few dm) can be present near top.
(0.05-0.4 m) Planar or undulated loaded bases (flames) and sharp to gradational tops. Lower to upper transitional plug flows rapidly decelerated when development of bands (sensu Baas et al., 2009) or turbulence-enhanced to lower transitional plug flows moderately decelerated when development of heterolithic tractional bedforms (Baas et al., 2016;Stevenson et al., 2020).
(0.5-1.2 m) Erosional (grooves) or sharp or loaded bases, and sharp tops or amalgamated. Planolites and Chondrites. High-density flows with high sediment fallout and suppressed traction (Kneller & Branney, 1995) in a moderately to well-oxygenated environment. Grooves indicate a phase with significant cohesive strength (Peakall et al., 2020).
Traction-plus-fallout from high-density flows and bedform aggradation in the upper-stage plane bed field with dominant unidirectional or oscillatory combined flow component (Tinterri, 2011). successions (thickness range of individual successions 10 to 40 m) (Fig. 6A) Leanza, 1990). Large bivalve shells (5 to 10 cm diameter) are either articulated and in life position or slightly disarticulated. Shells are associated with well-preserved circular crinoid ossicles (3 to 5 mm diameter) or articulated stems (2 cm long) often found in tuffaceous-rich levels, whereas rare siliceous sponge spicules are concentrated in micritic-rich levels. Strata are tabular with irregular, crenulated sharp bases and tops with some vertical borings (Trypanites). Stylolites, which can be bed-parallel or sub-parallel are also present, outlined by dark greyish insoluble material. Medium beds (50 to 70 cm thick) of fossiliferous carbonate (F1b) have a packstone texture that is massive or with rare undulose sub-parallel laminations (0.4 to 2.0 cm thick). Disarticulated bivalve shells in broken fragments (5 cm long) and very coarse sand-sized reworked volcaniclastic grains (quartz and K-feldspar) are scattered throughout beds; the bulk of the bed is composed of a fine-grained sand-sized fraction of this bioclastic and volcaniclastic material ( Fig. 8A). Beds are tabular with wavy sharp bases and tops, locally showing Pascichnia grazing traces (Helminthopsis) at the contact with the overlying marlstone (F3c) of FA3.

Interpretation
In the micritic tuffaceous carbonate facies (F1a), the preservation of bivalves in life position or weakly disarticulated indicates no or very little post-mortem transport. This indicates in situ deposition under stable low hydraulic conditions consistent with the small diameter of articulated crinoid stems. The abundance and species of large-sized bivalves indicate deposition in a well-oxygenated and low-energy shallow environment (cf. Damborenea & Manceñido, 1979. The in situ suspension-feeder organisms and carbonatesecreting benthos, including encrusters and shelly biocalcifiers (see also FA2), point to a warm-temperate carbonate factory (cf. Betzler et al., 1997;Fl€ ugel, 2004). The lack of early cementation and binding of these deposits favoured their reworking and resedimentation as allochemical biodetrital carbonate material (e.g. Halfar et al., 2004). Resedimented volcaniclastic grains together with parautochthonous broken or disarticulated shells in fossiliferous carbonate (F1b) suggest intermittent waveinduced physical reworking with limited transport, in an otherwise calm environment. This facies association is interpreted to represent warm-temperate carbonate platform deposits in a low-energy and well-oxygenated shallow environment below or near the fairweatherwave base, influenced by episodic ash-falls (e.g. Armella et al., 2016). Biodetrital carbonate mud accumulated in situ with skeletal accretion, bioerosion and weak mechanical reworking of organisms that formed a micritic substrate on fault-block highs, which remained near photic depths (<30 to 50 m) in normal marine waters. Primary volcanic material mixed with carbonate deposits indicates simultaneous explosive volcanic eruptions from the western magmatic arc which is postulated to have intermittently restricted the photic zone and depleted the biota (Armella et al., 2016).  (F2a) forms medium to thick beds (20 to 90 cm thick), with a very poorly sorted, very coarse to medium-grained, weakly to non-graded, sandy matrix. They (F2a) bear abundant polymict granule to gravel-sized (0.5 to 25 cm long) and angular to subrounded clasts (reworked pyroclastic and effusive volcanics, quartz and mudstone). Pebbly conglomerate deposits (F2a) also bear a few pieces of wood (up to 15 cm), disarticulated small bivalve and rhynchonellid shells (2 to 5 cm long), and pentagonal or circular crinoid ossicles (5 mm diameter). Individual beds are massive and tabular with low erosional relief at their bases (<10 cm deep), are locally clast-supported, and have sharp planar tops. Medium to thick beds (50 to 70 cm thick) of moderately sorted, planar or cross-laminated, crudely normally or inversely graded calcareous sandstone (F2b) are formed of bioclasts, reworked coarse sand-sized volcaniclastic grains (quartz, K-feldspar) and subrounded clasts (1.5 cm long) from a volcaniclastic protolith. Bioclasts include thick-walled fragmented mollusc shells (2 to 5 cm long), crinoid ossicles (5 mm diameter), fragments of discoidal robust solitary coral (<5 cm diameter) (Montlivaltia, see Gulisano & Guti errez Pleimling, 1995) and bryozoans (3 cm diameter) (Fig. 8B). Beds can show low-angle cross-laminations and may form isolated lenses, have a broad lenticular to lens-shape (<8 m wide) with erosional bases (up to 15 cm deep) and sharp planar tops, or are locally amalgamated. Medium beds (30 to 50 cm thick) of well-sorted, massive, calcarenite (F2c) have a grainstone texture, with well-rounded, medium to fine-grained skeletal and quartz grains and thin-shelled bivalves (1 to 2 cm long). Beds have sharp planar bases and tops, which locally are penetrated by vertical burrows (Skolithos and Ophiomorpha).

Interpretation
The pebbly conglomerate deposits (F2a) are characterized by a very poor sorting of immature reworked volcaniclastic clasts admixed with bioclasts, which indicate a weak gravitational sorting and reworking across a short transport distance prior to deposition. The dearth of mud matrix, local segregation of clasts at the base of beds, little to no grading and lack of stratification in these deposits suggest emplacement by subaqueous non-cohesive hyperconcentrated density flows (Mulder & Alexander, 2001;Drzewiecki & Sim o, 2002;Payros & Pujalte, 2008). Calciturbidites (F2b) contain a high proportion of fragmented bioclasts relative to reworked volcaniclastic material, with local normal or inverse grading and traction structures that indicate transport and deposition by high-density turbidity currents (Braga et al., 2001;Payros & Pujalte, 2008). Calcarenites (F2c) lack mud matrix, grading or structures and reflect intense bioclastic grain reworking and rounding prior to deposition by concentrated bioclastic grain flows (Drzewiecki & Sim o, 2002;Halfar et al., 2004). Given the composition and diversity of the bioclastic material in this facies association, it may have been sourced from erosion and reworking of sediment accumulated in the carbonate platform (FA1). Collectively, the grainy facies of FA2, which results from mechanical reworking of bioclastic and volcaniclastic material, formed platform-derived calciclastic aprons detached from the carbonate platform (FA1) across rift topography. Therefore, FA2 is interpreted to represent proximal periplatform deposits accumulated below the fairweather-wave base, in a deeper environment than FA1. The dearth of carbonate mud within FA2 suggests resuspension and elutriation of the biodetrital carbonate mud fraction of FA1, prior to deposition of the coarser-grained components. FA2 and the finergrained mud-rich FA3 do not interdigitate, rather FA3 overlies FA2, suggesting that during . Note that the datum is defined by the contact between Unit 2 and Unit 3 which corresponds to a facies change from F3b light grey laminated spiculitic marlstone (FA3) to F4a dark grey organic-rich calcareous mudstone with concretions and bivalve shell pavements (FA4). Facies are detailed in Table 1. deposition of FA2 there was efficient bypass of the resuspended mud fraction to deeper parts of the carbonate system (cf. P erez-L opez & P erez-Valera, 2012) represented by FA3.

Description
Facies Association 3 (FA3) corresponds to poorly stratified, medium to thin-bedded mudrich mixed carbonate-clastic successions (thickness range of individual successions 10 to 35 m thick) (Fig. 6B) including laterally extensive tuff layers (5 to 10 cm). FA3 unconformably overlies FA2, thickening towards fault-block lows and thinning on fault-block highs where it overlies FA1 (Fig. 7). Thin to medium beds (15 to 50 cm thick) of bioclastic carbonate (F3a) have packstone texture, with fine to medium sand-sized reworked volcaniclastic grains (quartz and Kfeldspar) from a volcaniclastic protolith, and small disarticulated shells of bivalves and brachiopods. Broken shells can be concentrated in undulose laminated normally graded grainrich layers (1 to 2 cm thick) (Fig. 8C). Individual beds have a tabular geometry with wavy sharp bases and tops that can show vertical burrows (Cylindrichnus) and can be rarely trough crossbedded. The medium beds (30 to 50 cm thick) of spiculitic carbonate (F3b) have a wackestone texture containing well-preserved monaxon and tetraxon siliceous sponge spicules (1 to 4 mm long) and a few silt to fine sand-sized reworked volcaniclastic grains (quartz and K-feldspar) from a volcaniclastic protolith. Pumice lapilli (1 to 2 mm diameter) and subrounded micritic intraclasts (2 to 5 cm long) are also found locally. Beds have an irregular tabular or lensshaped geometry with sharp bases and tops, and are locally bioturbated (Chondrites bollensis and Trichichnus). Locally, spiculitic carbonate beds are interbedded with medium beds (10 to 60 cm thick) of thinly laminated (0.5 to 5.0 cm thick) marlstone (F3c), including sponge spicules, tuffaceous and finely comminuted shell hash layers (0.5 cm thick) and rare large ammonites (15 cm diameter).

Interpretation
The normally graded and laminated grainy bioclastic carbonate beds (F3a), locally with rare trough-cross-bedding, record episodic high-energy storm-wave reworking (i.e. distal tempestite, P erez-L opez & P erez-Valera, 2012). The presence of Cylindrichnus in bioclastic carbonate beds indicates well-oxygenated sea-bottom conditions (Ekdale & Harding, 2015). In contrast, Chondrites and Trichichnus traces in spiculitic carbonate (F3b) and marlstone (F3c) record a decrease of oxygen levels at the sediment-water interface compared with the carbonate platform deposits (FA1 and FA2). The spiculitic carbonate beds bearing intraclasts, intercalated within marlstones, indicate reworking by low-energy storm-wave events alternating with periods of post-storm settling (i.e. stormwinnowed beds, P erez-L opez & P erez-Valera, 2012). In these fine-grained facies, the well-preserved siliceous sponge megascleres support limited transport and a parautochthonous origin, from a harder carbonate substrate below the storm-wave base. Allochemical bioclastic material and biodetrital carbonate mud in FA3 might have initiated with storm-wave reworking of unconsolidated carbonate platform substrates on the highs (FA1) and efficient exportation of the resuspended mud fraction towards more distal parts of the carbonate system (cf. P erez-L opez & P erez-Valera, 2012). Frequent dilution by volcanic influxes is indicated by pumice and tuffaceous material (D'Atri et al., 1999;Halfar et al., 2004). The lower biota diversity in FA3 (compared to FA1 and FA2) dominated by allochthonous to parautochthonous disarticulated bivalves, brachiopods and siliceous sponges supports deposition at greater water depths than FA1 and FA2. This facies association represents distal periplatform deposits emplaced under moderate to low-energy conditions, near or below the storm-wave base in an offshore-transition environment, with the progressive deepening and reduction from welloxygenated to moderately-oxygenated sea-bottom conditions. The stacking of distal periplatform deposits (FA3) on proximal periplatform deposits (FA2) in fault-block lows, and carbonate platform deposits (FA1) on fault-block highs, might record the retrogradation of the carbonate system.

Facies Association 4: Sand-starved basin
Description Facies Association 4 (FA4) forms poorly stratified, very thin to thin-bedded calcareous mudstone-dominated successions (thickness range of individual successions 20 to 120 m thick), which unconformably overlie FA3 and are overlain by FA5 (Fig. 6D). FA4 is locally affected by significant thickness changes across rift structures (up to 100 m across a few kilometres, see Fig. 5). Calcareous mudstone (F4a) is massive to faintly laminated, with well-preserved dispersed carbonaceous matter, and is mainly composed of silt-size and clay-size carbonate material. Very thin to thin-bedding is monotonous (1 to 10 cm and up to 20 cm thick), commonly deflected by oblate calcareous strata-bound concretions (<15 cm long) and intercalated with thin tuffs (1 to 5 cm thick). Pyrite is present disseminated within discontinuous layers (<0.5 cm thick) parallel to bedding or within elliptical oblong concretions (5 to 8 cm long) (Fig. 8D). Faunal content is represented by small ammonites (2 to 5 cm diameter) and articulated or disarticulated shells of juvenile bivalves (0.5 to 2.0 cm long) distributed along bedding planes, as shell pavements [Posidonotis cancellata (Leanza)]. Some thin to medium-beds (5 to 40 cm thick) of massive to graded calcareous mud-rich siltstone (F4b), and rare graded medium to fine-grained bioclastic calcareous sandstone (F4c) with lowangle planar and current-ripple lamination, can be intercalated within mudstone. Bioclastic sandstone beds contain abundant broken skeletal material and have sharp planar or erosional bases with tool marks, and sharp tops (Fig. 8E). Bioturbation in calcareous mud-rich siltstone (F4b) includes small forms of Chondrites intricatus and Phycosiphon traces.

Interpretation
The thin-bedding, massive character and wellpreserved carbonaceous matter in calcareous mudstone (F4a) suggest in situ pelagic biogenic production by calcareous organisms, which is well-established in the study area during the Early Jurassic (mainly coccolithophorids, Angelozzi & P erez Panera, 2016). Allochthonous biodetrital carbonate mud plumes exported from platform environments by storm-wave-induced offshore transport (Schieber, 2016;Birgenheier & Moore, 2018) might also have contributed to accumulation of calcareous mudstone. Bioclastic calcareous sandstone (F4c) and siltstone (F4b) with normal grading, current ripples and low angle-lamination, in the lower part of FA4, suggest clastic dilution by storm-wave enhanced, bioclastic-rich, low-density turbidity currents (Table 1; Bouma, 1962;Lowe, 1982). The high organic matter content of type II recorded elsewhere in the Lower Los Molles Formation (determined with measurements of organic carbon isotopes and hydrogen index) and with TOC between 2% and 11%, supports a mixture of marine and terrestrial components (cf. Al-Suwaidi et al., 2016). This is consistent with a mixed in situ and allochthonous source for mud and organic matter. Preservation of organic matter was favoured by the very limited bioturbation and clastic dilution in this low-energy and low-oxygen environment with low sedimentation rates (cf. Birgenheier & Moore, 2018). Reducing conditions below the seabed promoted pyrite mineralization and pre-compaction seabed diagenetic processes that formed the calcareous concretions, which record very low sedimentation rates or breaks in sedimentation (Taylor et al., 1995). The pavement-type concentration of juvenile, low-oxygen tolerant bivalve specimens [Posidonotis cancellata (Leanza), cf. Damborenea et al., 2013] (see Fig. 2) records episodes of high mortality events and/or condensed surfaces with very low sedimentation rates. Little to no post-mortem bottom current reworking of bivalves, and scarcity of silty and sandy beds, support deposition well below the storm-wave base, consistent with estimations of palaeobathymetry at the base of the Los Molles Formation ranging between 50 to 100 m and 200 to 400 m (cf. G omez Omil et al., 2002;G omez-P erez, 2003). In summary, this facies association records deposition of pelagic material and allochthonous fine-grained carbonate sediments exported by storms from the coeval drowned carbonate platforms. These sediments deposited in a sandstarved basinal environment with prevailing poorly oxygenated sea-bottom conditions. for the location of sections). The stratigraphic increase in bed thickness and grain-size, amalgamation rate and increasing sand content upward in Unit 4 suggest progradation of the lobe complex. Note that lobes pass systematically down-dip and laterally (from section 5 to 9 see Fig. 10) into thinner-bedded (tens of centimetres thick) HEB-rich sandy heterolithic margins (HEB types 2 and 3), which pinch-out farther than the abrupt pinch-out of metres thick debrites (HEB type 1). Note that sections 1 and 2 show a lateral facies change in the dirty lobes whereas sections 5 to 9 show a downdip facies change, according to their dominant north-east palaeocurrent direction (see current directions in Fig. 5). In the cleaner lobes, sections 5 to 9 show a more oblique to downdip facies change, supported by the dominant north-west orientation of sole marks, and a consistent north-east direction recorded by current ripples and cross-bedding (see current directions in Fig. 5). Fig. 10. Panoramic view from uncrewed aerial vehicle (UAV) photograph (cars on the road for scale) showing the onlap limit of the Lower Los Molles Formation (Unit 3) onto the Chachil Formation (Units 1 and 2), the location of the compaction hinge and location of dirty and cleaner lobes, which form the intraslope lobe complex (ca 5 km minimum wide 9 6 to 8 km long 9 50 to 70 m thick) within Unit 4 (colours for lobes are specified in Fig. 9.). Respective lateral and frontal pinch-outs are indicated. The clastic sill body (1.5 to 3.8 m thick, 5 to 8 km across) (see Figs 5 and 9) steps 1 to 2 km outward from the onlap margins of the lobe complex. (A) View of slumped mudstone and siltstone interval (1.6 m thick and kilometre-scale). (B) View of the clastic sill injectite that splits into smaller sills with abrupt pinch-out terminations across <2 km on the El Luchador fault-block high. Note that the deformed zone is an area affected by inverse faults.

Interpretation
Massive to subtly graded mudstone (F5.1a) suggests deposition from waning fine-grained dilute muddy turbidity currents (Stow & Bowen, 1980). The pin-striped laminated mudstone (F5.1b) was deposited by clay-laden, turbulence-enhanced transitional flows, to lower transitional plug flows (sensu Baas et al., 2009), rapidly decelerating with possible shear sorting and mixing that formed thin clay-bearing and silt to sandbearing stripes (cf. 'streaky bedding' of Baas et al., 2016). Graded structured siltstone and crudely to normally graded sandstone beds (F5.1c and F5.1d) were deposited with tractional reworking by high to low-density turbidity currents allowing differential particle settling (Lowe, 1982;Best & Bridge, 1992). Sparse bioturbation and dissolution of calcitic shells preserved as moulds indicates deposition under more oxygenated sea-bottom conditions than the mudstone deposits of FA4. This can be explained as in the oxic zone, decomposition and remineralization of organic matter produces carbon dioxide, which in turn forms carbonic acid that can lead to dissolution of calcium carbonate shells (Aller, 1982;Morse et al., 2007). The discordant massive sandstone bodies (F5.1e) lacking any grading or sedimentary structure, with conspicuous cementation and planar layers of sub-spheroidal vugs, are interpreted as clastic injectites including minor, thin dykes associated with laterally extensive thick sills (Hurst et al., 2011). Lack of clay matrix, the presence of subangular clasts, and 'in situ' rafts, indicate the incorporation of lithified host strata during injection of slow moving laminar flows, and the distribution of large mudstone clasts mantling surfaces of sills might result from entrainment and abrasion by erosive injecting flows (Cobain et al., 2015). Muddy heterolithic successions (FA5.1) with narrow grainsize range, beds with gradational tops and tabular extensive geometry for hundreds of metres, lack of amalgamation, local erosion with few decimetre-scale sandstone-prone scours and sparse bioturbation suggest deposition in a poorly to moderately-oxygenated distal lobe fringe setting (Mutti, 1977;Pr elat & Hodgson, 2013).

Facies Association 5.2: Intraslope lobes -Lobe fringe
Description Facies Association 5.2 (FA5.2) comprises thin to medium-bedded sandy heterolithic successions (<5 to 25 m thick) (Fig. 6F), including some thick beds, extends for up to 5 km and is transitional to FA5.1, vertically and laterally (Fig. 9). Beds are rarely amalgamated, and include very poorly sorted sandy mudstone and muddy sandstone facies with variable mud matrix and clast content. Bed geometry is irregular, with common pinch and swell and lateral facies changes with abrupt thinning and pinch-out across less than tens of metres (Figs 9 and 10). The distinct facies divisions recognized in individual beds of FA5.2 suggest that they are mainly hybrid event beds (HEBs; sensu Haughton et al., 2009) and three main bed types are identified (Fig. 12).
The chaotic muddy sandstone (F5.2e) (4 to 6 m thick) is characterized by a very poorly sorted, patchy medium to fine-grained sand-rich matrix bearing outsized granules, coarse sand grains and abundant mud chips. Beds have erosional bases and mounded tops (Fig. 11B). The matrix supports pebble to cobble-sized mudstone and sandstone clasts (5 to 30 cm diameter), deformed sand-streaks (30 cm long) and heterolithic rafts (up to 80 cm long), with mud-rich and clast-poor tops. The chaotic sandy mudstone division (F5.2f) (0.8 to 5.0 m thick) has a very poorly sorted sandy mudstone matrix containing outsized coarse sand grains, plant material and pebble-sized subangular to subrounded mudstone, siltstone and sandstone clasts (5 to 50 cm long), locally including well-preserved shallow-marine ostreid shells (Fig. 11C). The matrix content in these beds increases upward where the largest clasts are segregated, floating in the matrix (Fig. 11C). Locally, the chaotic division (F5.2e or F5.2f) is overlain by a normally graded sandstone with planar to undulose laminations and/or ripples (F5.1d) (<0.2 m thick), or by a clast-rich muddy sandstone (F5.2c) (20 to 60 cm thick) with irregular pinch and swell bed geometry, often mud-filled, and a sheared erosive basal contact (shown in Fig. 6F).

Interpretation
Type 1 HEBs (Fig. 12) are interpreted to reflect deposition from a thick forerunning debris-flow associated with development of a basal concentrated density flow as a result of shear mixing at flow-interfaces (with surrounding seawater) during a single flow event (Amy et al., 2005). In these HEBs, the chaotic muddy sandstone (F5.2e) and sandy mudstone (F5.2f) were likely deposited by intermediate to high yield strength debrisflows (sensu Talling et al., 2012), with significant entrainment of compacted substrate (Dakin et al., 2013;Talling, 2013). Entrainment of ambient water might have diluted the basal part of the flow enabling substrate erosion and/or hydroplaning (Marr et al., 2001). Fluid mixing at the base of the debris-flow likely played a significant role in decreasing the debris-flow strength below the point at which it could form a rigid plug flow. This resulted in a basal layer (F5.2a) that behaved as a transitional flow (Baas et al., 2011), where shearing and breaking of clay particles were likely sufficiently high, and cohesive bed yield strength sufficiently low, to enable the progressive deposition of sand. Locally, low-amplitude bed-waves (cf. Baas et al., 2016;Baker & Baas, 2020) developed at the top of the basal sand layer beneath slowly decelerated lower to The facies in brackets are only locally present. HEB type 1 comprises a basal massive mud-poor sandstone with mud-rich bed-waves with low-angle laminae at top (F5.2a) (see inset picture) and pinch and swell geometry, sharply overlain and locally scoured by a chaotic muddy sandstone (F5.2e) or sandy mudstone (F5.2f). This is in turn overlain by a clast-rich muddy sandstone (F5.2c) with pinch and swell geometry and erosive basal contact, and capped by massive silty mudstone (F5.2d) filling topography at top. HEB type 2 comprises a basal massive to laminated mud-poor sandstone (F5.2a) that grades into planar sub-parallel to undulose banded muddy sandstone with mudchips (F5.2b) that can be absent and eroded by a clast-rich muddy sandstone (F5.2c) capped by massive silty mudstone (F5.2d). HEB type 3 comprises a massive mud-poor sandstone (F5.2a) that grades into a well-developed banded muddy sandstone (F5.2b) locally with mudstone clasts (centimetre-scale) at base and heterolithic bedforms, capped by massive silty mudstone (F5.2d). The banded sandstone can have elongated mudstone clasts (2 to 5 cm long) at the base, and includes planar sub-parallel to undulose and low-angle draping sandy and muddy laminae associated with symmetrical hummock-like heterolithic bedforms and/or asymmetrical large current ripples with high-angle foresets. Note the difference in clast lithology and size between thick HEB (type 1) and thin HEBs (types 2 and 3). upper transitional plug flow (sensu Baas et al., 2009). In places, this basal sand was eroded to give the pinch and swell geometries that were later infilled with mud trapped by the topography. The development of the basal sandy layer and thus the occurrence of mud-rich bedwaves (F5.2a), as well as variations in the nature of the overlying bed, might be strongly related to lateral changes in debris-flow strength and the irregular erosional behaviour of the debris-flow itself (Talling, 2013). The upper sandy divisions (F5.1d or F5.2c) overlying the debrite might have formed through dilution and shear mixing at the top and front of the debris-flows (Talling et al., 2002;Mohrig & Marr, 2003;Felix et al., 2009). This led either to formation of a turbulent cloud that evolved into a low-density turbidity current (F5.1d), or to increased concentration as a result of mixing with the underlying debris-flow muddy material and evolution into a transient low to intermediate yield strength sandy debris-flow (F5.2c) (Talling et al., 2012).
In type 2 HEBs, the locally developed sub-parallel to undulose banded division (F5.2b) was emplaced beneath rapidly decelerating lower to upper transitional plug flows (Baas et al., 2016;Stevenson et al., 2020). Where present the banded division is frequently cut by a debritic division emplaced by a more cohesive, intermediate (F5.2c) to high (F5.2f) yield strength debris-flow (Talling et al., 2012). In type 3 HEBs, the banded division is well-developed, also including heterolithic tractional bedforms emplaced beneath moderately decelerated turbulence-enhanced transitional flows, to lower transitional plug flows (Baas et al., 2016;Baker & Baas, 2020;Stevenson et al., 2020). The overlying massive silty mudstone cap (F5.2d) records consolidation after deposition of cohesive silty fluid mud flows (Baas et al., 2011). The deposition of HEBs 2 and 3 is linked to thinner decimetre-scale facies divisions in comparison to HEB 1 (Fig. 12). Their development might result from flow bulking through entrainment of clayey substrate (Talling et al., 2004), deceleration and flow transformation of an initial highdensity turbidity current into, lower or upper transitional plug flow (HEB 3), and quasi-laminar plug flow or laminar plug flow (type 2 HEBs) (Baas et al., 2011(Baas et al., , 2016Peakall et al., 2020). Note that a forerunning turbidity current (i.e. Haughton et al., 2009) can be excluded in the formation of these type 1 and 2 HEBs as the basal sand layers pinch out prior to the overlying debritic divisions. Sandy heterolithic successions (FA5.2) show complicated spatial facies relationships and bed thickness changes associated with pinch and swell geometry, common metre-scale mud-filled scours and hybrid event bed development. These characteristics support a lobe fringe subenvironment interpretation (Hodgson, 2009) but with more variable bed thickness pattern, including thick beds, with more erosion, flow transformation and HEBs present in both frontal and lateral lobe fringes (sensu Spychala et al., 2017), and associated with the development of heterolithic bedforms (Baker & Baas, 2020;Stevenson et al., 2020).

Facies Association 5.3: Intraslope lobes -Lobe axis
Description Facies Association 5.3 (FA5.3) forms crudely to well-stratified, medium to thick-bedded sandstone-dominated successions (5 to 12 m thick) ( Fig. 6G and H), which can include some HEBs and transitionally overlie or pass downdip into FA5.2 (Fig. 9). Tabular extensive sandstone packages (1.5 to 5.0 m thick) are characterized by amalgamated bed contacts with up to 20 cm relief marked by abrupt grain-size breaks and mudstone clasts, or comprise thin mudstone interbeds (5 to 15 cm thick). The upper parts of sandstone beds are locally intensely burrowed by Planolites and Chondrites.
Massive to crudely stratified sandstones (F5.3a) form medium to thick beds (0.5 to 1.2 m thick), which are poorly sorted, mud-poor (i.e. low argillaceous matrix compared to FA5.1 and FA5.2 facies), coarse to medium-grained and locally structured with diffuse planar parallel laminations (1 to 3 cm thick). Beds are often amalgamated, with sharp planar or erosional bases displaying groove casts and bearing elongated discrete subrounded lithic granule to pebble-sized mudstone clasts (1 to 6 cm long). When not amalgamated, the upper part of beds can grade crudely into banded muddy sandstone (F5.2b) and/or laminated fine-grained sandstone (F5.1d).
Granular sandstone (F5.3b) forms thin to medium beds (5 to 90 cm thick), which are very poorly sorted, very coarse-grained or coarse to medium-grained, structureless and weakly normally graded into structured medium-grained sandstone (F5.3c). Locally, it can contain abundant subangular granule-sized grains (0.2 to 0.4 cm diameter) and pebble-size mudstone and siltstone clasts (5 to 8 cm long), forming either an inverse grading pattern near the bed base or a coarse-tail grading pattern throughout beds. Some mudstone clasts armoured with a mix of quartz pebbles and bioclasts, including fragments of belemnites, bivalve shells and planktonic foraminifera (possibly Globigerinids), can be found in these beds (Fig. 11G). Beds can be isolated with erosional bases and sharp tops, or amalgamated. Locally, they can stack into crossbedded sets (1.5 to 2.0 m thick, up to 5 m wide) dipping up to 10° (Fig. 11E).

Interpretation
Massive to crudely stratified coarse to mediumgrained sandstones (F5.3a) were deposited with high sediment fallout rates that could suppress traction (Kneller & Branney, 1995). Lower fallout rates and collapse of high concentration nearbed laminar sheared layers enabled the formation of diffuse laminae (Sumner et al., 2008). Massive granular sandstone with crude normal or inverse to normal coarse-tail grading (F5.3b) suggests near-bed transport of highly concentrated coarse particles with bedload traction and deposition by high-density to hyperconcentrated density flows (Lowe, 1982;Mulder & Alexander, 2001). Structured medium-grained sandstone (F5.3c) was deposited by traction-plus-fallout beneath stratified high-density combined flows, with high sediment fallout rates, but enabling bedform aggradation in the upper-stage plane bed stability field and bedform migration with bedload traction (Tinterri, 2011). Asymmetrical rounded ripples formed with a dominant unidirectional flow component, whereas symmetrical hummocky bedforms might have developed with a dominant oscillatory combined flow component (Tinterri, 2011). The oscillatory flow component could not originate with surface waves given the absence of wave-induced or storm-induced structures and the low diversity of bioturbation traces in FA5, which point to deposition from sediment gravity flows below the storm-wave base. Alternatively, these bedforms are interpreted to form from flow deflection against a confining slope, and resulting interactions of reflected internal wave trains with the near-bed unidirectional flow component (cf. Tinterri, 2011). The effects of confinement and basin configuration are discussed in the section about Unit 4.
Medium to thick-bedded sandstone-dominated successions (FA5.3) comprise amalgamated sandy packages (1.5 to 5.0 m thick, ca 1 to 2 km across) of continuous tabular beds with basal erosion (<0.6 m deep) and locally narrow and shallow incisional features interpreted as sandstone-prone scour-fills (1.5 to 2.0 m deep, ca 5 m long), suggesting a lobe axis sub-environment (Etienne et al., 2012;Pr elat & Hodgson, 2013). The increase in bioturbation intensity compared with FA5.1 indicates deposition in a moderately to well-oxygenated environment, consistent with frequent extrabasinal siliciclastic influxes.

SYNTHESIS OF DEPOSITIONAL SYSTEMS AND ARCHITECTURE OF TECTONO-SEDIMENTARY UNITS
Genetic facies relationships described in the previous section and stratal relationships across structures described in this section allowed four tectono-sedimentary units to be defined (Figs 3 and 5). Analysis of the spatial facies distribution, palaeocurrents, thickness changes, stratal bounding surfaces and associated angular relationships between units are used to decipher the late synrift to early post-rift evolution of sedimentary systems during the Early Jurassic (Figs 2 and 13). Units 1 and 2 belong to the Chachil Formation, and Units 3 and 4 correspond to the Lower Los Molles Formation Present-day mean bedding dipdirections have been calculated for each unit in the three different geographic zones (Fig. 4) in order to show the differences in bed orientation and dip, and unconformities, between the different units and structural domains. In each zone, there is a stratigraphic decrease in stratal dip angle and from a multidirectional dip-direction pattern to a dominant east/south-east dip azimuth trend (Fig. 4). The terms 'proximal' and 'distal' refer to the position relative to the Southern Chachil Graben margin (Fig. 3A).

Unit 1
Unit 1 includes carbonate platform deposits (FA1) that onlap onto fault-block highs with an angular discordance between ca 4°and 14°and proximal periplatform deposits (FA2) that onlap onto fault-block lows with an angular discordance between ca 11°and 15° (Fig. 5). Carbonate platform deposits (FA1) have moderate stratal dip angles (ca 16 to 25°) with an east/south-east dip direction on the Morro del Aguila (Zone SW) and Puesto Alfaro (Zone Central) faultblock highs, and a south-west dip direction on the El Luchador fault-block (Zone NE) (Fig. 4). Proximal periplatform deposits (FA2) show moderate stratal dip angles (ca 20°) with an east dip direction in the Mirador de Chachil fault-block low (Zone SW) and higher dip angles (ca 33°) with a south dip direction in the Pic un Leuf u fault-block low (Zone Central) (Fig. 4).
The fault-block carbonate platforms (FA1) are detached from the basin margin (sensu Bosence, 2005) (Figs 7 and 13A), and exhibit subtle internal growth stratal patterns with pinch-out towards the crest of rotated fault-block highs in the immediate hangingwall, downdip of the Southern Chachil Graben margin (Figs 5 and 7). The Morro del Aguila and the Paine Milla faultblocks host extensive carbonate platforms (2 to 5 km length and width, up to 40 m thick) dominated by micritic tuffaceous carbonate (F1a). The platform, nucleated on the Puesto Alfaro fault-block highs, is dominated by fossiliferous laminated carbonate (F1b) (Fig. 7). The fauna of Unit 1 records the flourishing of a suspension feeder-dominated carbonate-secreting benthos including subtidal encrusters (scleractinian corals, bryozoans and crinoids) and shelly biocalcifiers (bivalves and brachiopods), which ensured productivity of the warm-temperate carbonate system (cf. Fl€ ugel, 2004). Although Unit 1 might record the downdip reworking of higher-energy inner ramp deposits, the limited exposure does not enable the reconstruction of the 3D geometry and complex facies distribution of the carbonate system. On the Southern Chachil Graben margin, a tide-influenced inner carbonate platform with tidal flats and subtidal ponds developed, and recorded shallowing-upward cycles with progradation of intertidal and supratidal deposits onto subtidal deposits (Armella et al., 2016). Carbonate sedimentation on the horst which formed the Southern Chachil Graben margin (Fig. 3) was strongly affected by ashfall and episodic subaerial exposure associated with firmground development in peritidal conditions, oxide coated surfaces and Glossifungites ichnofacies (Armella et al., 2016). Nonetheless, these surfaces associated with subaerial exposure have not been observed on the smaller fault-block highs within the Chachil Graben, suggesting that they remained submerged (Fig. 3). Therefore, a combination of subaerial degradation on the Southern Chachil Graben margin and enhanced current erosion of faultblock highs might have controlled extensive reworking of the poorly consolidated carbonate platform deposits (FA1) and volcano-sedimentary substrate into proximal periplatform deposits (FA2).
Mixed carbonate-clastic proximal periplatform deposits (FA2) accumulated in the Mirador de Chachil and Pic un Leuf u fault-block lows with onlap onto volcano-sedimentary syn-rift deposits (Lapa Formation) and pinch-out towards intervening fault-block highs (Figs 7 and 13A). The polymict composition of pebbly conglomerates (F2a), dominated by brecciated reworked volcaniclastic material, were deposited by hyperconcentrated density flows that originated from small-scale failures (cf. Drzewiecki & Sim o, 2002) triggered by fault-block tilting and destabilization of primary volcanic deposits (Lapa Formation) (Fig. 13A). Similar deposits identified as basal conglomerates of the Chachil Formation have been recognized in the subsurface in the south-east of the Neuqu en Basin, and are interpreted as evidence for subaerial exposure and degradation of the Precuyano Cycle deposits (Schiuma & Llamb ıas, 2008). Calciturbidite deposits (F2b) can be compared to carbonate platform-derived calciclastic aprons lacking internal organization and with local development of basal lags in small channel-fills (cf. Braga et al., 2001;Payros & Pujalte, 2008). The stratigraphic increase in proportion and sorting of grainy allochemical carbonate material, sourced from carbonate platforms on fault-block highs, recorded throughout FA2 might reflect the decreased availability of the volcano-sedimentary syn-rift substrate for reworking due to progressive onlap into fault-block lows, together with a decrease of destabilization events with fault-block tilting.
South-south-east oriented palaeocurrents (Fig. 5), stratal fanning of proximal periplatform deposits (FA2) in the Mirador de Chachil fault-block low (5 to 25 m thick across 3 to 4 km) towards the Morro del Aguila fault-block high, and relative sea-level fall and subaerial exposure of the Southern Chachil Graben margin, suggest active faulting and southeastward tilt of the Mirador de Chachil fault-block low. Facies distribution, depositional geometries and variability of stratal dip azimuths across structures (Figs 3 and 4) reflect active fault-block rotation during deposition of Unit 1 and support its late syn-rift development (Cross & Bosence, 2008;Dorobek, 2008).

Unit 2
Unit 2 corresponds to mud-rich mixed carbonate-clastic distal periplatform deposits (FA3) that form wedge-shaped packages in fault-block lows, thin drapes on fault-block highs, and onlap onto Unit 1 to form a ca 3°to 10°angular discordance (Fig. 5). In the Mirador de Chachil proximal fault-block low (Zone SW) and Pic un Leuf u distal fault-block low (Zone Central), Unit 2 presents moderate stratal dip angles (ca 18 to 32°) and consistent east/south-east dip direction. This contrasts with the higher dip angles (ca 62°) and the north-east dip direction of Unit 2 across the El Luchador fault-block (Zone NE), which record the overprint of later inversion tectonics (Fig. 4).
Bioclastic carbonate (F3a) deposited in the Mirador de Chachil proximal fault-block low shows thinning (20 to 12 m thick across 3 to 4 km) and onlap onto calciturbidite (F2b) towards the Puesto Alfaro fault-block high (Figs  5 and 7). These deposits are overlain by interbedded spiculitic carbonate and marlstone (F3b-F3c) that thin (15 to 5 m thick) towards the Puesto Alfaro fault-block high. Bioclastic carbonate (F3a) also thins (18 to 8 m thick across <2 km) from the El Luchador fault-block high towards the Pic un Leuf u fault-block low (Figs 7 and 13B). Overlying interbedded spiculitic carbonate and marlstone (F3b and F3c) thin (from 10 to 5 m thick) from the Pic un Leuf u fault-block low towards the El Luchador faultblock high (Figs 7 and 13B). On the Morro del Aguila, Puesto Alfaro and Paine Milla faultblock highs, bioclastic and spiculitic carbonate (F3a and F3b) are absent and carbonate platform deposits (F1a and F1b) are draped by marlstone (F3c) (5 to 10 m thick) (Figs 7 and 13B). This contact is marked by Pascichnia grazing traces at the top of carbonate platform deposits representing a condensed surface on the fault-block highs due to sedimentation under poorly oxygenated and relatively deep bottom water conditions (Ekdale & Mason, 1988). A similar situation is observed on the Southern Chachil Graben margin where the carbonate platform deposits are overlain by marlstone (e.g. Armella et al., 2016).
The stratigraphic deepening recorded by Unit 2 provides strong evidence for fault-controlled subsidence, given that tectonically induced relative sea-level changes could outpace the lowamplitude and rates of eustatic rise during Early Jurassic greenhouse time (cf. Ravn as & Steel, 1998). Deepening and reworking of unconsolidated carbonate platforms (FA1) on the most elevated fault-block highs bounding the graben, promoted redeposition of allochemical fine biodetrital carbonate material (FA3) in faultblock lows (e.g. Halfar et al., 2004). The thinning-upward and fining-upward trend of Unit 2 deposits in the Mirador de Chachil and Pic un Leuf u fault-block lows, and thickening towards the Morro del Aguila and Puesto Alfaro faultblock highs dominated by condensed sedimentation, support syn-depositional relative sea-level rise and differential subsidence (Fig. 13B). The stratal pattern suggests an increased displacement along the Chihuido Bayo fault system bounding the Chachil Graben and localized normal faulting in the main hangingwall with formation of the El Luchador fault-block high (Fig. 13B).
Tectonic subsidence and relative sea-level rise outpaced carbonate sedimentation rates and controlled retrogradation of the carbonate system. Drowning of the carbonate system culminated with establishment of a deep-marine environment and reduction of oxygen recorded near the top of Unit 2 from the Southern Chachil Graben margin to the interior of the Chachil Graben.
The deterioration of chemical and physical conditions with tectonically-induced relative sealevel rise contributed to drowning of the carbonate factory across the entire Chachil Graben (e.g. Santantonio, 1994;Ruiz-Ortiz et al., 2004;Navarro et al., 2012).

Unit 3
Unit 3 forms a calcareous mudstone-dominated succession (FA4), which overlies Unit 2 locally with an unconformable relationship, with a ca 3°to 9°bedding difference between Unit 2 and Unit 3 at the graben margins. Dip angles are moderate (ca 16°to 25°) with general south-east dip direction in the Mirador de Chachil proximal fault-block low (Zone SW) and Pic un Leuf u distal fault-block low (Zone Central). This contrasts with higher dip angles (ca 57°) with some east/south-east dip direction measured in the deformed strata affected by later inversion tectonics across the El Luchador fault-block (Zone NE) (Fig. 4).
The base of Unit 3 infills the intrabasinal topography inherited at the top of Unit 2 burying the Puesto Alfaro fault-block highs, and therefore records a change of basin geometry with infilling of the inherited rift topography. This is supported by the thinning of Unit 3 (<20 m thick) to the north-east across the El Luchador and Paine Milla fault-block highs and to the south-west across the Morro del Aguila, and by the thickening (70 to 120 m thick) from the Mirador de Chachil to Pic un Leuf u faultblock low (Fig. 5). The deformed steeply dipping strata of Unit 3 near the El Luchador faultblock (Zone NE) is associated with abrupt mudstone thickening (up to 100 m offset of compacted strata) across less than 2 km from the Paine Milla platform towards the adjacent Pic un Leuf u fault-block low (Figs 5 and 13C). This thickness change occurred across north-west/ south-east striking and south-west dipping rift faults involved in the formation of the El Luchador fault-block (Fig. 13B). The potential process that caused small-scale deformation recorded in Unit 3 is detailed in the Discussion.
The occurrence of Posidonotis cancellata (Leanza) shell pavements in the lower part of Los Molles Formation has been reported in TOC-rich mudstone of another marine rift depocentre, located 20 km south-east of the Chachil Graben (Al-Suwaidi et al., 2016). Al-Suwaidi et al. (2016) report negative carbon isotope excursions associated with the Toarcian Oceanic Anoxic Event (TOAE) during the late Tenuicostatum-early Dactylioceras Hoelderi Andean Ammonite Biozone (AAB 15 to 16) (Fig. 2). The TOAE might be recorded in Unit 3, which spans the latest Pliensbachian-Early Toarcian in the Chachil Graben, marked by a reduction of benthic fauna diversity recorded from Unit 2 to Unit 3 (Fig. 2) and variable sedimentation rates related to important storm activity under a warm temperate climate (cf. Volkheimer et al., 2008). These conditions prevailed over south-western Gondwana, and were not favourable for the deposition of organic black shales, which characterize the TOAE in the Northern Hemisphere, where a warm and humid climate favoured the development of anoxia (Dera & Donnadieu, 2012;Fantasia et al., 2018). The stratigraphic evolution from carbonate to terrigenous mudstone recorded at the top of Unit 3 (Fig. 6D) marks the transition to Unit 4 and indicates an increase of fluvio-deltaic runoff, which could be related to climatic change towards more humid conditions from the early Late Toarcian (cf. Volkheimer et al., 2008).

Unit 4
Unit 4 represents a submarine lobe complex (sensu Pr elat et al., 2009) whose base is transitional with, or sub-concordant to Unit 3, except around the El Luchador fault-block (Zone NE) where it shows an angular discordance between ca <1 to 8°onto Unit 3. Steeper dips in Zone NE, with bedding differences of up to 36°b etween Unit 3 and 4 (Fig. 4), are related to structures that formed during later inversion tectonics, but these are highly localized and overall the depositional architecture remains well-preserved. Dip angles of Unit 4 are moderate (ca 16°to 21°) with consistent east/south-east dip direction (Fig. 4). Unit 4 heals across structures and shows general thickness changes with thickening (150 to 170 m thick) from the Mirador de Chachil to the Pic un Leuf u fault-block low and thinning (50 to 100 m) across the Morro del Aguila, El Luchador and Paine Milla fault-block highs (Figs 5 and 9).
The lower part of Unit 4, dominated by muddy heterolithic strata, corresponds to distal lobe fringes (FA5.1) with a low sandstone proportion (Figs 9 and 10). The upper part of Unit 4 is characterized by a higher proportion of sandstone and includes stacked sandy heterolithic strata in lobe fringes (FA5.2), sandstone-dominated strata in lobe axes (FA5.3) and minor muddy heterolithic strata in distal lobe fringes (FA5.1) (Fig. 9). Individual lobes have a low aspect ratio (1.5 to 5.0 m thick, few kilometres across) and collectively these deposits form a 50 to 70 m thick, ca 5 km minimum wide and 6 to 8 km long intraslope lobe complex (Figs 9 and  10). The intraslope lobe complex comprises basal dirty lobes (Figs 9 and 12) mainly represented by lobe fringe deposits with a high proportion of HEBs (types 1, 2 and 3) and locally by lobe axis deposits (Fig. 6G). The dirty lobe deposits are characterized by a finer grain-size, argillaceous matrix and high clast content. The overlying cleaner lobes include lobe axis deposits (Fig. 6H), locally with thick HEBs (type 1), and lobe fringe deposits with thin HEBs (mainly types 2 and 3) (Figs 9 and 12). The cleaner lobe deposits are characterized by thicker beds, coarser grain-size and lower argillaceous matrix and clast content (Figs 9 and 14).
The intraslope lobe complex within Unit 4 records a rapid increase of extrabasinal siliciclastic supply, favouring oxygenation and increase of bioturbation intensity and diversity (Figs 2 and 13D), consistent with the warming humid climate (cf. Volkheimer et al., 2008). The increase in volume and frequency of sand-rich sediment gravity flows over time, associated with the stratigraphic increase in bed thickness and grain-size, amalgamation rate, and increasing sand content upward in the lobes of Unit 4 suggest progradation of the lobe complex (Crabaugh & Steel, 2004;Macdonald et al., 2011). The slightly shingled pattern of lobe sub-environments, high thickness and facies variability across short distances (100 m) (Fig. 9), palaeocurrent evidence for flow reflection and deflection (Fig. 5), and widespread erosion and scouring in proximal and distal lobe fringe settings (Figs 6E, 6F and 11D) may result from the development of lobes in a partially confined setting (Figs 10 and 13D).
Partial confinement could have been controlled by inherited topography across the El Luchador fault-block high, which formed a south/south-west-facing oblique counterslope that could have influenced sediment gravity flow behaviour and led to deviation of the general north-east flow direction mainly towards the south-east and rarely south-west (Fig. 5). This could explain the opposing directions recorded in current-ripple laminations (Fig. 8F) and combined-flow bedforms (Figs 11F and 12), showing azimuth dispersion up to 180°and sole marks (mainly grooves) with average azimuth dispersion of 45°(see palaeocurrents, Fig. 5). Interaction of flows with relief is suggested by the development of a range of combined-flow bedforms (Fig. 11F) (cf. Tinterri, 2011) and heterolithic bedforms (Fig. 12), which could indicate seabed topography (cf. Hofstra et al., 2018). Therefore, these bedforms with multiple palaeoflow directions might have developed with oblique flow reflection and deflection (i.e. Kneller et al., 1991;Amy et al., 2004) against the counterslope flanking the El Luchador faultblock high (Figs 13D and 14).
Partial confinement is further supported by the abrupt thinning of thick-bedded amalgamated sandstone bedsets (1.5 to 5.0 m thick) across a few kilometres, without marked grainsize or sorting change and by the frontal pinchout pattern of the lobe complex. The lobes pass systematically down-dip into thinner-bedded (tens of centimetres thick) HEB-rich sandy heterolithic succession (HEB types 2 and 3), which pinch-out farther than the abrupt pinch-out of metres thick debrites (HEB type 1) (Figs 9 and 12) (see Discussion). Block diagram showing the evolution of sedimentation patterns and interactions with topography in the Chachil Graben from an underfilled to a sediment-balanced depocentre during the syn-rift to post-rift transition, detailing relationships between facies distribution and structures. Detailed logs in lobes show typical facies association of different lobe sub-environments in both dirty and cleaner lobes. Lobe axis includes medium to thickbedded sandstone-dominated successions with amalgamated packages (1.5 to 5.0 m thick) and can comprise HEBs, which form intralobe fluid flow baffles. Lobe fringe includes thin to medium-bedded sandy heterolithic successions dominated by HEBs. Distal fringe includes thin-bedded muddy heterolithic successions dominated by pin-striped mudstone and sandstone with common streaky bedding (i.e. banding). Intraslope lobes are characterized by significant erosion in axis and fringe, and well-developed tractional banded/heterolithic bedforms and combined flow bedforms and thin across a few kilometres, without marked grain-size or sorting change. Note the frontal pinch-out of the lobe complex into thinner-bedded (tens of centimetres thick) HEB-rich sandy heterolithic succession dominated by HEB types 2 and 3, which pinch-out farther than the abrupt pinch-out of metres thick debrites of HEBs type 1.
The presence of the sill-dominated injectite network stratigraphically above, below and lateral to the lobe complex, and their pinch-out towards the Paine Milla fault-block high, supports the influence of inherited topography. The relief that induced substantial confinement of lobes would also have later promoted overpressure and clastic injection (e.g. Cobain et al., 2017). Laterally extensive sills (1.5 to 3.8 m thick, 5 to 8 km across) step 1 to 2 km outward from the onlap margins of the lobe complex and split with abrupt pinch-out terminations across the El Luchador fault-block high (Fig. 10). The linear ridges present on the surface of some sills show a general north-east/ south-west orientation which supports a northwest/south-east direction of crack propagation, approximately parallel to the south/south-westfacing counterslope (e.g. Kane, 2010). This together with the offset pinch-out between lobes and sills at this location might reflect the influence of the buried El Luchador fault-block high on the morphology of the injectite complex (e.g. Cobain et al., 2017).
The overpressure prior to injection may have been a combination of: (i) deep-seated hot basinal fluid and gas expulsion through rift faults related to a mature hydrocarbon plumbing system in the syn-rift deposits and/or sourced from crustal magmatic activity; and (ii) differential loading and compaction of buried strata (e.g. Boehm & Moore, 2002;Kane, 2010). Widespread expulsion of basinal fluids has been reported locally with the formation of methane seepages associated with development of bioherms at the seabed during deposition of the Lower Los Molles Formation (G omez-P erez, 2003), which is equivalent to Unit 3 of this study. This would be consistent with a primary mechanism of overpressure build-up. Migration of basinal fluids into the shallowly buried but tilted sandstone bodies of Unit 4 may have triggered injection (e.g. Cobain et al., 2017), although a seismic trigger for injection cannot be ruled out.

Early post-rift inherited topography
The analysis of the syn-rift to post-rift transition in the Chachil Graben has revealed major thickness changes within the organic-rich mudstone (Unit 3) across the Paine Milla-El Luchador fault-block high, and the role of topography recorded by intraslope lobe architecture (Unit 4). The local and small-scale character of deformation in early post-rift strata of the Lower Los Molles Formation, which has not been observed elsewhere in other adjacent depocentres at a regional-scale, is consistent with the end of riftrelated extension at that time in the study area. Here, the processes and significance of this deformation are discussed.

Unit 3
Local thickness change (>100 m of compacted strata) at the transition between the Paine Milla-El Luchador fault-blocks and the adjacent Pic un Leuf u fault-block low must have been accommodated across a major buried/blind rift structure during deposition of Unit 3 (Figs 3A and 5). The limited facies and thickness changes in Unit 2 across this fault do not support a large bathymetric difference between the El Luchador faultblock high and the adjacent Pic un Leuf u faultblock low at this time (Figs 7 and 13B). Therefore, additional differential subsidence across this fault occurred after deposition of Unit 2 to create accommodation in the Pic un Leuf u faultblock low associated with thickening of the mudstone-dominated succession of Unit 3 (Figs 5 and 13C). The presence of syn-depositional topography across the fault during deposition of Unit 3 is further supported by the facies change from calcareous mudstone with interbedded siltstone and bioclastic sandstone beds in the Pic un Leuf u fault-block low, to condensed homogeneous mudstone on the El Luchador and Paine Milla fault-block highs. The presence of multiple tuff beds, which act as markers, in Unit 3 (and more rarely in Unit 4) on fault-block highs and lows support that these facies changes in mudstones were stratigraphically equivalent. The sedimentation rates from compacted strata (cf. McArthur et al., 2013) are estimated at 4.3 m/Myr in Unit 3 considering the maximum sediment thickness on the structurally elevated fault-block highs, and 27.6 m/Myr considering the maximum sediment thickness in the axis of the Pic un Leuf u fault-block low. These rates are based on measured thicknesses of compacted strata (Fig. 5), taking into account the minimum thickness (ca 20 m thick) on fault-block highs and maximum thickness (ca 127 m thick) in the deepest fault-block low (Pic un Leuf u). The duration of deposition of Unit 3 of ca 4.6 Myr is based on the biostratigraphy (using the Posidonotis cancellata biozone at the base of Unit 3), and U-Pb ages at the base (184.4 AE 2 Ma) and top (179.8 AE 1.4 Ma) of Unit 3 (Privat, 2019) (see Fig. 2). Whilst sedimentation rates are estimates, they are in the range of pelagic sedimentation rates similarly calculated from compacted mudstone strata in the Kimmeridge Clay Formation (20 to 45 m/Myr, cf. McArthur et al., 2013), which comparably formed an extensive, organic-rich mudstone blanket across rift topography in the North Sea. Moreover, stratal geometries of Unit 3 and Unit 4 record the long-lived influence of this fault-block relief, which is postulated to have formed a compaction hinge above the inferred blind fault tip that controlled deep-water sediment dispersal during deposition of Unit 4 (Figs 13D and 14). If syn-depositional relief was formed with riftrelated extensional tectonics, one might expect near-fault damage (for example, fracturing and folding) and widespread normal faulting in other parts of the Chachil Graben and in adjacent depocentres, which is not the case.
This study examines whether this syn-depositional relief which developed during deposition of Unit 3 could result from differential compaction of the buried early syn-rift strata of the Lapa Formation across the pre-existing rift fault bounding the Pic un Leuf u fault-block low. At this location, the Lapa Formation comprises a thick volcano-sedimentary succession, based on the thickness change recorded towards the north-east from the Puesto Alfaro (up to 200 m thick) to the Piletas section (up to 500 to 600 m thick) (cf. Franzese et al., 2006). The compaction of a thicker syn-rift succession in the Pic un Leuf u fault-block low would result in a larger compaction-induced accommodation in this location (cf. Carminati & Santantonio, 2005). This difference in accommodation explains the accumulation of a thicker mudstone succession (120 m thick) than the stratigraphically equivalent and thinner succession (20 m thick) in the Puesto Alfaro and El Luchador fault-block highs. Changes in both lithology and thickness of the syn-rift succession (Lapa Formation), recorded from the Puesto Alfaro faultblock high to the Piletas fault-block low (see fig. 5 in Franzese et al., 2006), might have enhanced the differential compaction and subsidence in the Pic un Leuf u fault-block low.
Although detailed backstripping of the entire syn-rift to post-rift succession is not possible in this study, based on the lack of constraints on a number of parameters and limitations of 3D exposure (see the Data and methodology section), it is possible to provide an estimate of compaction. This in turn supports the postulated role of differential compaction in explaining early post-rift subsidence patterns. In the Pic un Leuf u fault-block low, it is difficult to quantify the exact thickness of the Lapa Formation because whilst the upper part is exposed, the lower part is lying in the subsurface as it dips and thickens towards the north-east, and is better exposed 2 km northward in the Piletas section (up to 500 to 600 m thick; Franzese et al., 2006). The Piletas section is used as an analogue for the Pic un Leuf u section, and therefore a thickness of 500 m is assumed, including an estimate of the most deeply buried lower half of the Lapa Formation (250 m thick) lying in the subsurface. The Lapa Formation comprises interbedded compactable clastic rocks and less compactable volcanic rocks (mostly matrix-supported breccias and minor conglomerates, sandstone and mudstone; Franzese et al., 2006), which can comprise up to 40% of the total thickness of the syn-rift succession (cf. Franzese et al., 2006). Considering the sandy matrix of breccias and conglomerates, it is assumed that their porosity is commensurate with sandstone which is typically deposited with 40 to 50% porosity that is reduced up to 15 to 25% at ca 1 to 2 km of burial depth (cf. Sclater & Christie, 1980;Skuce, 1996). Therefore, lithostatic loading of the upper half of the Lapa Formation and the Chachil Formation (as a total of 280 m thick) could compact up to 20% of the original thickness of the lower half of the Lapa Formation The original thickness of the lower Lapa Formation at time of deposition (ca 313 m thick) means that mechanical compaction would have provided ca 63 m of accommodation space in the Pic un Leuf u fault-block low.
Similarly, the mudstone succession at the base of the Lower Los Molles Formation (Unit 3), could compact up to 40% of its original thickness, as mudstone is deposited with 70 to 80% porosity and undergoes a ca 40% porosity loss at ca 200 m of burial depth (cf. Shinn & Robbin, 1983;Skuce, 1996). The decompacted thickness of mudstone would have been ca 200 m in the Pic un Leuf u fault-block low and ca 100 m on the Puesto Alfaro fault-block high. Therefore, the ca 100 m thickening of Unit 3 recorded in the Pic un Leuf u fault-block low could have filled the accommodation space (ca 63 m) generated with mechanical compaction of the lower Lapa and Chachil Formation and healed the inherited fault-scarp relief. This would be consistent with the thickness (ca 20 m) of compacted mudstone of Unit 3 found on the El Luchador and Paine Milla fault-block highs.
The amount of compaction-induced accommodation proposed here is likely underestimated by using conservative compaction rates, as the sandstone compaction rate applied to the Lapa Formation does not account for the higher compaction rate of interlayered mudstone in this mixed succession. The extra accommodation which would have been generated with mechanical compaction of deposits of the Chachil Formation is also neglected. However, these approximations highlight that the thickness changes recorded in the mudstone succession of Unit 3 in the Pic un Leuf u fault-block low, are in the range of the amount of accommodation that could be created with differential compaction of buried strata across this fault-scarp relief.
Surface deformation, including small-scale faulting and large-scale folding, can result from differential compaction across irregular rigid basement topography and/or arise due to lithological heterogeneity (e.g. Barr, 1991;Skuce, 1996;Cristallini et al., 2006Cristallini et al., , 2009Lohr & Underhill, 2015). Compaction-driven surface deformation proceeds over a relatively short timescale (i.e. a few millions of years) and might have outpaced the low sedimentation rate during deposition of Unit 3. This increased the fault scarp relief and the gradient of the hangingwall in the Pic un Leuf u fault-block low, which exceeded the critical sediment stability threshold (5°) for carbonate muds (cf. Kenter, 1990). Gravitational remobilization and downslope redistribution of fine-grained sediment with slumping occurred, sometimes as metres-scale contorted beds (Fig. 10). Therefore, the presence of this relief is supported by the occurrence of syn-sedimentary destabilization events during deposition of Unit 3, which progressively healed the fault-scarp relief (i.e. 'parallel onlap fill', Cartwright, 1991). As the compaction and surface deformation rate outpaced the very low sediment accumulation rate, the fault-scarp relief increased 'passively', enhanced by differential compaction and subsidence, without the need to invoke active extension and fault slip (e.g. Carminati & Santantonio, 2005).
These observations have implications for the interpretation of the syn-rift to post-rift transition and indicate that the timing of active (i.e. syn-rift) normal faulting in the Lower Los Molles Formation could be erroneously inferred based on relatively subtle changes in thickness and facies, which instead could be controlled by post-rift differential compaction. Additionally, these results support findings in other locations of the Neuqu en Basin (north-east and south-east of the Huincul High), where small-scale normal faulting, stratal wedging and graben-scale folding of Early Jurassic strata around rift structures are similarly related to differential compaction of the buried syn-rift succession (Cristallini et al., 2006(Cristallini et al., , 2009.

Unit 4
Unit 4 is characterized by thinning with onlap updip towards the NNW-SSE trending Southern Chachil Graben margin, thickening and dipping towards the east/north-east into the graben. This configuration defines a broad (ca 6 km across), NNE-trending, long-wavelength and low-amplitude monocline slightly oblique to the Southern Chachil Graben margin (Figs 1C and 3A). The gentle dip and lack of uplifted hangingwall strata with major thrust deformation in Unit 4 near the Southern Chachil Graben margin suggest that the monocline did not grow in response to compression during basin inversion. Instead, monocline folding could have developed during the early post-rift in response to differential compaction across the Southern Chachil Graben margin and enhanced accommodation in the centre of the graben, as documented in other rift basins (e.g. Barr, 1991;Cartwright, 1991;Skuce, 1996;Monaldi et al., 2008;L opez-Gamund ı & Barragan, 2012;Lohr & Underhill, 2015). Therefore, early post-rift strata record the effects of differential compaction at: (i) a local scale in Unit 3 with the formation of a compaction hinge above a buried rift structure in the graben; and (ii) a larger scale in Unit 4 with the development of a long-wavelength and lowamplitude monocline fold along the Southern Chachil Graben margin (Figs 3A and 10). Constraining the effects of compaction-induced deformation across rift structures is key to estimating the timing of the syn-rift to post-rift transition at depocentre-scale, as well as for predicting early post-rift reservoir distribution with combined stratigraphic and structural traps.

Controls on sedimentation during the syn-rift to post-rift transition
The transition from the late syn-rift to early post-rift is marked by a change in stratigraphic architecture, sediment dispersal (i.e. from axial to transverse to structural strike) and from intrabasinal carbonate to extrabasinal siliciclastic sediment supply. Stratal stacking patterns in the Chachil Graben record a change from general retrogradation to progradation, and from underfilled to sediment-balanced conditions (Figs 2,  13 and 14). The relative importance of rift basin topography, tectonics and climate as key controls on these changes are discussed herein.
Late syn-rift Linkage of the Chachil Graben bounding fault system during the rift climax (Franzese et al., 2006), and the presence of acidic volcanic and granitic rocks resistant to weathering along the Southern Chachil Graben margin might have limited the development of local drainage catchments bounding the depocentre during the late syn-rift (Leeder et al., 1998). The relatively low relief inherited from the overfilled depocentre configuration (sensu Muravchik et al., 2011) also limited intrabasinal clastic sediment supply and reworking of volcano-sedimentary syn-rift deposits (Lapa Formation). In contrast, steeper gradients preserved in the adjacent Cat an-Lil Graben promoted intense fault-block degradation and accumulation of late syn-rift fan deltas (Muravchik et al., 2014).
Growth of normal faults during the late synrift segmented the immediate hangingwall of the Southern Chachil Graben margin and induced a structural pattern that differed from the rift climax structural configuration (cf. Franzese et al., 2006). These conditions, together with the temperate climate (cf. Volkheimer et al., 2008) that was characterized by reduced precipitation, erosion, and both intrabasinal and extrabasinal sediment delivery, permitted the late syn-rift development of the carbonate system in the Chachil Graben (Fig. 14). Fault linkage physically disconnected footwall-derived sediment sources from their sink and subsequent marine transgression associated with rapid subsidence suppressed sediment input by reducing the size of the intrabasinal drainage catchments that were flooded. Thus, instead of acting as an efficient clastic intrabasinal source during the late synrift and early post-rift, the horst structure which formed the Southern Chachil Graben margin underwent rapid marine flooding and acted as a submerged barrier that limited clastic supply. Fault-block highs controlled the dimensions and morphology of shallow-marine isolated carbonate platforms (Fig. 14). Syn-rift deformation prevented the development of fault-block carbonate platforms in down-faulted depressions or their coalescence into a larger composite platform, which would have required tectonic quiescence and subdued fault-related relief (Dorobek, 2008). Instead, fault-block lows received proximal periplatform mixed carbonate-clastic deposits reworked from the carbonate platforms on faultblock highs, which were funnelled across the complicated structural rift pattern, mainly axially to the structural strike towards the southeast (Fig. 14). With tectonic subsidence and rising relative sea-level, the carbonate platform supplied greater flux of fine-grained reworked carbonate sediment basinward (e.g. Dorobek, 2008). This resulted in retrogradation of the system and accumulation of distal periplatform mixed carbonate-clastic mud-rich deposits into fault-block lows, whereas condensed sedimentation occurred over carbonate platform deposits on fault-block highs (Fig. 14). The warm-temperate nature of the carbonate system (cf. climate Volkheimer et al., 2008) played a key role in its organization and evolution across rift structures. The low energy setting and lack of early cementation and binding (no framebuilding biota) of these carbonate deposits favoured their reworking into allochemical biodetrital carbonate material towards fault-block lows. Furthermore, their low production rates were insufficient to keep pace with tectonically-driven deepening of water depths, which led to rapid drowning. This might explain why these warm-temperate systems remain poorly documented in rift settings (Gulf of California, Halfar et al., 2004) compared to their tropical counterparts (Gulf of Aden and Suez Rift, Cross & Bosence, 2008). The late syn-rift carbonate sedimentation, and drowning of the system in response to increased rates of fault-controlled subsidence, meant that the Chachil Graben evolved into an underfilled depocentre, which had implications for the subsequent development of the early post-rift system (Fig. 14).
Early post-rift During the early post-rift, potential local intrabasinal fault-block sources were rapidly submerged with the back-arc evolution of marine rift depocentres involving high regional subsidence rates related to rapid thermal subsidence effects (cf. Cristallini et al., 2009). Regional marine flooding and mud-draping of inherited rift topography resulted in a very limited to absent intrabasinal sediment supply from degradation of local fault-block highs. Sand starvation prevailed in marine rift depocentres until the onset of sediment supply from extrabasinal sources along basin margins (Fig. 1A). In contrast to rapid thermal subsidence and regional drowning of fault-blocks, rift basins associated with passive margins often result in an early post-rift configuration where local subaqueous or subaerial degradation of fault-block highs lasts longer due to slow thermal subsidence and provides intrabasinal clastic supply (e.g. Zachariah et al., 2009;Jarsve et al., 2014;Henstra et al., 2016).
In the Chachil Graben, the development of early post-rift lobes with an extrabasinal sediment source is suggested by mud-draping of intrabasinal rift topography, palaeocurrents with respect to rift structures and sandstone composition, which does not match local granitic bedrock on footwall blocks. The early post-rift lobe sandstones are characterized by a high amount of terrestrial organic material (wood, leaves and carbonaceous fragments), the presence of ostreid-bearing clasts found in some of the mudrich debrites (Fig. 11C) and armoured mudstone clasts (Fig. 11G). These elements support a main provenance from hinterland sources, and shelf to slope sediments (Hodgson, 2009;Migeon et al., 2010;Talling et al., 2010) for the lobe deposits in the Lower Los Molles Formation. An extrabasinal source with high sediment yield would have been required to develop the coarsening-upward intraslope lobe succession that prograded north-eastward into the Chachil Graben. Palaeoflows recorded by early post-rift lobes were mainly transverse to the structural strike of rift fault-blocks, suggesting that most of the former complicated rift topography in the graben was healed with mudstone-dominated and heterolithic-dominated successions (Fig. 5). Therefore, clastic sediment supply, which caused the Chachil Graben to evolve into a sediment-balanced depocentre during the early postrift, played a major role in infilling the inherited rift topography (Figs 2 and 14). This is consistent with the warm humid climate that prevailed from the early Late Toarcian (cf. Volkheimer et al., 2008) and promoted conditions suitable for high amounts of erosion and riverine runoff, delivering high volumes of sediment towards the basin despite rising sea-level and overall highstand (e.g. Leeder et al., 1998;Yu et al., 2013;Bal azs et al., 2017).
During the early post-rift, extrabasinal sediment supply can be trapped by inherited rift topography in proximal depocentres, which are progressively filled allowing coarse sediment to bypass downslope to more distal depocentres (e.g. Lien, 2005;Soares et al., 2012). At a regional scale, during the early post-rift, the inherited rift topography preserved in the south-western Neuqu en Basin (Legarreta & Uliana, 1996;Burgess et al., 2000;G omez Omil et al., 2002) would have promoted trapping of clastic material along the nascent shelf-slope system located >20 to 25 km south-east of the Chachil Graben, close to the hinterland source (see Fig. 1A). This basin configuration and semi-arid to temperate conditions at the onset of the early post-rift (cf. Fig. 2) might have inhibited the coeval supply of extrabasinal siliciclastic sediment reaching distal depocentres of the southern shelf-slope system. Consequently, deposition of thick organic-rich calcareous mud occurred across the deeply [up to 400 m, cf. G omez Omil et al. (2002), G omez-P erez (2003)] submerged inherited rift topography of distal depocentres, such as the Chachil Graben, which remained sand-starved. This promoted the longlived preservation of major rift basin topography, such as the horst structure which formed the Southern Chachil Graben margin and enhanced local accommodation, with differential compaction and subsidence across buried rift faults into the graben (Fig. 14).

Implications for the characteristics of intraslope lobes in early post-rift settings
The depositional architecture, facies distribution, termination style and related HEBs of the early post-rift intraslope lobes in the Lower Los Molles Formation (Figs 9, 12 and 14) provide rare insights into the characteristics of these particular deep-water lobes enabling comparison with other basin-floor or base-of-slope lobes (e.g. Haughton et al., 2003;Southern et al., 2017;Dodd et al., 2019).

Relationship between inherited relief and intraslope lobe characteristics
The architecture of the documented intraslope lobe complex records a stratigraphic change from dirty to cleaner lobe facies, with an increase in grain-size, bed thickness and amalgamation, and sand matrix and clast content, associated with progradation of the intraslope lobe system (Figs 9 and 10). A potential mechanism explaining the stratigraphic change from dirty to cleaner lobes could be the progressive healing of relief obstructing the flow pathway, which would reduce flow interactions with the seabed topography and limit flow transformation. However, evidence of palaeoflow deflection and reflection, combined-flow and heterolithic bedforms (hummock-like bedforms, and asymmetrical rounded ripples) in both dirty and cleaner lobes support sustained interactions of flows with intrabasinal topography. This relief was likely generated by the compaction hinge, which acted as an oblique counterslope (Fig. 14) for flows and promoted the northeastward progradation of the intraslope lobe complex.
Hybrid event beds (HEBs) are developed throughout the lobe complex stratigraphy, with lobe complex pinch-outs being dominated by the thin type 2 and 3 HEBs that pinch-out farther basinward than the thicker, type 1 HEBs. Therefore, the presence of frontal topography is interpreted to have prevented systematic segregation of thick HEBs (type 1) towards the fringe of the lobe complex. This pattern of HEB distribution contrasts with documented onlaps in other basins, where the HEBs are only present close to the pinch-out of lobe complex, suggesting that flow transformation occurs near the topography (e.g. Pyles & Jennette, 2009;Patacci et al., 2014).
The sedimentological characteristics and architecture of early post-rift intraslope lobes (Figs 9 and 12) differ from unconfined to weakly confined base-of-slope and basin-floor lobe models. Typically, the latter are larger (>10 km width and length), with more gradual thinning rates and transitions between lobe sub-environments, well-defined lobe off-axes, erosion and scouring restricted to lobe axes, and thinner HEBs (centimetre to decimetre-scale) (Pr elat et al., 2009;Pr elat & Hodgson, 2013;Spychala et al., 2017). The early post-rift intraslope lobe complex in the Chachil Graben has a low aspect ratio (ca 5 km minimum wide 9 6-8 km long 9 50-70 m thick) and well-defined pinchout terminations similar to previously documented intraslope lobes (6-10 km wide 9 15-25 km long 9 10-15 m thick, Spychala et al., 2015; 8 km wide 9 8 km long 9 120 m thick, Jobe et al., 2017). The documented lobes show scouring and lobe-scale compensation patterns, which are also reported in confined intraslope lobes characterized by sub-seismic-scale heterogeneity (cf. Jobe et al., 2017). However, they are significantly coarser-grained and more argillaceous, with HEB-rich lobe fringes that contrast with sandier fringes of other intraslope lobes, which are HEB-poor Jobe et al., 2017). The characteristics of the documented lobes in the Los Molles Formation which could be representative of early post-rift intraslope lobes that developed on mud-draped seabed topography are: abrupt transitions between lobe sub-environments (<100 m), widespread erosion including in the lobe fringe, HEBs in axial and fringe sub-environments of dirty and cleaner lobes, and combined-flow and heterolithic bedforms (Figs 9  and 12).

Hybrid event bed type origin
The poorly sorted, relatively coarse-grained and matrix-rich beds in the intraslope lobes implies deposition from flows that transported coarsegrained material with a high clay and silt component, which reduced the settling velocity of large grains (cf. Al-Ja'aidi et al., 2004). The heterogeneity recorded in the dirty lobes, which mark the first stage of development of the intraslope lobe complex, is characterized by an abundance of relatively thin HEBs (types 2 and 3) with similar thicknesses of facies divisions (Fig. 12). Their deposition may result from autogenic processes  with flow transformation through entrainment of the widespread muddy substrate after an extended period of sand-starvation, which increased flow cohesion, and enhanced flow stratification and collapse (Baas et al., 2011;Kane et al., 2017). In contrast, thicker HEBs (type 1) containing a well-developed debritic division encased by sandy divisions (Fig. 12) are common both in the dirty and cleaner lobes, and comprise large clasts of various lithologies that record substrate erosion. The sporadic distribution of these HEBs, which results from flow transformation with shear mixing at top and base of voluminous, long run-out debris-flows transporting shallow-marine bioclasts, seems to be the product of larger flows sourced from mass-wasting processes in a proximal shelf-slope setting (e.g. Hodgson, 2009;Migeon et al., 2010;Talling et al., 2010).
Despite both allogenic and autogenic controls on the type of HEBs, the stratigraphic decrease of intrabed-scale heterogeneity could reflect common HEB development during periods of initiation and growth of out-of-grade systems (Haughton et al., 2003;Hodgson, 2009). Therefore, the reduction of thin HEBs (types 2 and 3) in the thicker-bedded and coarser-grained cleaner lobes (Figs 9 and 14) could be associated with progradation of the intraslope lobe system and maturation of sediment routing pathways, which together reduced the entrainment of muddy substrate across topography.
The distinctive characteristics of intraslope lobes developed in the Lower Los Molles Formation shows that more investigations of outcrop and subsurface analogues are required to develop models for the architecture and evolution of intraslope lobes encountered in early post-rift settings.

CONCLUSIONS
This study advances understanding of the organization of sedimentary systems with physiographic transformation of rift topography, change in sedimentation regime, and sediment routing and supply during the syn-rift to postrift transition in the Chachil Graben. The late syn-rift evolution of the Chachil Graben as an underfilled depocentre, with intrabasinal carbonate sedimentation and mainly axial sediment routing, was controlled by differential tectonic subsidence and rapid deepening, which led to subsequent retrogradation and drowning of the carbonate system. Later draping of the inherited rift topography by organic-rich calcareous mud inhibited intrabasinal sediment supply sources and recorded onset of the early post-rift with local and small-scale deformation associated with differential compaction and subsidence. Abrupt increases in extrabasinal clastic supply enabled the healing of the former rift topography and evolution of the Chachil Graben into a sediment-balanced depocentre, with the progradation of an intraslope lobe complex mainly transverse to buried rift structures. This study highlights the influence of an early post-rift compaction hinge formed across a rift fault in the graben, on local thickness changes within organic-rich mudstone at the base of the Lower Los Molles Formation and on sedimentological characteristics and architecture of intraslope lobes. The intraslope lobe complex (ca 5 km minimum wide 9 6-8 km long 9 50-70 m thick) is associated with a kilometre-scale silldominated clastic injectite network that steps towards the compaction hinge. The intraslope lobe complex records a stratigraphic evolution from dirty to cleaner lobes, with a decrease of intralobe bed-scale heterogeneity related to the distribution of hybrid event beds (HEBs). As the system prograded north-eastward, the compaction hinge acted as an oblique counterslope which induced partial confinement and interacted with a range of sediment gravity flows. This resulted in the development of combinedflow and heterolithic bedforms, widespread erosion, scouring and HEB distribution in both axis and fringe sub-environments of dirty and cleaner lobes, with rapid lateral transitions of lobe sub-environments, and HEB-rich pinch-out style of the intraslope complex.
These results support the need to better constrain the effects of differential compaction across rift structures, which can be misinterpreted as originating from active tectonics, in order to estimate the timing of the syn-rift to post-rift transition at depocentre-scale. The documented intraslope lobe complex shows the risk for reservoir deterioration (i.e. HEBinduced lobe heterogeneity) and post-burial remobilization (i.e. injectite complex), which have implications for the prediction of reservoir distribution in combined stratigraphic and structural traps.
paper, which helped to improve Fig. 10. Associate Editor Jim Hendry, Telm Bover Arnal and an anonymous reviewer are thanked for their detailed and constructive reviews that significantly improved the paper. We also thank Editor Peir Pufahl.