Ice‐marginal forced regressive deltas in glacial lake basins: geomorphology, facies variability and large‐scale depositional architecture

This study presents a synthesis of the geomorphology, facies variability and depositional architecture of ice‐marginal deltas affected by rapid lake‐level change. The integration of digital elevation models, outcrop, borehole, ground‐penetrating radar and high‐resolution shear‐wave seismic data allows for a comprehensive analysis of these delta systems and provides information about the distinct types of deltaic facies and geometries generated under different lake‐level trends. The exposed delta sediments record mainly the phase of maximum lake level and subsequent lake drainage. The stair‐stepped profiles of the delta systems reflect the progressive basinward lobe deposition during forced regression when the lakes successively drained. Depending on the rate and magnitude of lake‐level fall, fan‐shaped, lobate or more digitate tongue‐like delta morphologies developed. Deposits of the stair‐stepped transgressive delta bodies are buried, downlapped and onlapped by the younger forced regressive deposits. The delta styles comprise both Gilbert‐type deltas and shoal‐water deltas. The sedimentary facies of the steep Gilbert‐type delta foresets include a wide range of gravity‐flow deposits. Delta deposits of the forced‐regressive phase are commonly dominated by coarse‐grained debrisflow deposits, indicating strong upslope erosion and cannibalization of older delta deposits. Deposits of supercritical turbidity currents are particularly common in sand‐rich Gilbert‐type deltas that formed during slow rises in lake level and during highstands. Foreset beds consist typically of laterally and vertically stacked deposits of antidunes and cyclic steps. The trigger mechanisms for these supercritical turbidity currents were both hyperpycnal meltwater flows and slope‐failure events. Shoal‐water deltas formed at low water depths during both low rates of lake‐level rise and forced regression. Deposition occurred from tractional flows. Transgressive mouthbars form laterally extensive sand‐rich delta bodies with a digitate, multi‐tongue morphology. In contrast, forced regressive gravelly shoal‐water deltas show a high dispersion of flow directions and form laterally overlapping delta lobes. Deformation structures in the forced‐regressive ice‐marginal deltas are mainly extensional features, including normal faults, small graben or half‐graben structures and shear‐deformation bands, which are related to gravitational delta tectonics, postglacial faulting during glacial‐isostatic adjustment, and crestal collapse above salt domes. A neotectonic component cannot be ruled out in some cases.

Ice-marginal deltas are excellent palaeogeographical archives, recording the glaciation history of marine and continental basins. These deltas commonly evolve from ice-contact systems to glacifluvial deltas during icemargin stillstand and retreat (Lønne 1995;Dietrich et al. 2017) and delta foreset-topset contacts can be used as water-level indicators if shoreline features are poorly developed or became eroded by later peri-and paraglacial processes (Winsemann et al. 2009Perkins & Brennand 2015;Lang et al. 2018). In remote areas delta morphology and the dimensions of feeder channels can be used as an important record of palaeo-lakes and the magnitude of surface-water flow (Martin & Jansson 2011;Villiers et al. 2013).
The depositional architecture of delta systems is a sensitive archive of short-and long-term base-level changes and many delta studies during the last 20 years focussed on a sequence-stratigraphic interpretation of marine systems and their response to global sealevel change (Postma 1995;Posamentier & Morris 2000; Uli cn y 2001; Catuneanu et al. 2011). However, there is still a need for a better understanding of the facies variability, progradation styles and large-scale depositional architecture of forced-regressive ice-marginal depositional systems, which are less well understood compared to non-glacigenic sedimentary environments (Brookfield & Martini 1999;Powell & Cooper 2002;Gutsell et al. 2004;Hirst 2012;Lang et al. 2012;Nutz et al. 2015;Dietrich et al. 2017;Gilbert et al. 2017).
This study presents a synthesis of the geomorphology, facies variability and large-scale depositional architecture of ice-marginal deltas controlled by rapid lake-level change. The selected field examples are considered to be representative of delta styles in glacial lake basins. The integration of digital elevation models, outcrop, borehole, ground-penetrating radar (GPR) and highresolution shear-wave seismic data allow assessment of the role of rapid base-level change in delta morphology, of deltaic facies and geometries generated under different lake-level trends, helping to recognize a hidden record of such changes where deltaic systems are poorly exposed.

Delta styles and depositional processes
Gilbert-type ice-marginal deltas commonly reflect a relatively stable position of the ice-margin in front of mountain ranges or bedrock highs that acted as pinning points (Powell 1990;Ashley 1995;Lønne 1995;Winsemann et al. 2007Winsemann et al. , 2011Girard et al. 2015). If the ice terminus remains stable for a longer period of time, grounding line fans may also aggrade to lake level and form an ice-contact/glacifluvial delta (Powell 1990;Lønne 1995). Shoal-water ice-marginal deltas may form in low-gradient settings along lake-basin strike, during lake-level rise on drowned Gilbert-type delta plains, or during lake-level fall (Ashley 1995;Winsemann et al. 2009;Eilertsen et al. 2011). The sediment supply is dominated by ephemeral meltwater flows from glaciers and massflows from hill slopes whereby the sediment yield from hill slopes strongly depends upon the local availability of (melt)water. In glaciolacustrine environments, the sediment-laden meltwater is typically denser than the surrounding lakewater and the deposition on delta slopes is therefore likely to be dominated by a wide range of gravity flows, with comparatively minor inputs from high-level suspended sediment (Ashley 1995;Lønne & Nemec 2004;Winsemann et al. 2011).
The complex morphology and depositional architecture of delta systems are the result of an interplay of water discharge, sediment supply and available accommodation space (Dunne & Hempton 1984;Postma 1995;Posamentier & Morris 2000;Muto & Steel 2001, 2004; Uli cn y 2001; Lønne & Nemec 2004;Ritchie et al. 2004a, b;Petter & Muto 2008;Eilertsen et al. 2011;Winsemann et al. 2011;Gobo et al. 2014Gobo et al. , 2015. The geological setting and type of dam exert key controls on proglacial lake growth and drainage. Major controlling factors are the location of the ice margin, elevation and topography of the surrounding landscape and the location and elevation of the lake-overflow channel(s) (Teller 1987;Kehew & Teller 1994;Carrivick & Tweed 2013;Lang et al. 2018). In contrast to glaciomarine settings, glacial lake basins are typically characterized by an initial baselevel rise during glacier advance, as the glacier blocks drainage outlets. Forced regression characteristically occurs during deglaciation when lake outlets are opened and rapid lake-level falls may occur (Kehew & Teller 1994;Ashley 1995;Brookfield & Martini 1999;Winsemann et al. 2011;Carrivick & Tweed 2013;Winsemann et al. 2016).
The depositional architecture of delta systems is therefore a sensitive archive of short-and long-term baselevel changes (Fig. 1), with the formation of delta-brink rising trajectories during base-level rise and delta-brink subhorizontal or falling trajectories during base-level stillstand and fall (Posamentier & Morris 2000;Catuneanu et al. 2011;Gobo et al. 2015). During transgression, the high rates of lake-level rise, which are common in glacial lake basins (Oviatt et al. 1992;Winsemann et al. 2011), will cause a rapid landward shift of delta-front lobes (Posamentier & Morris 2000;Catuneanu et al. 2011) and the formation of a stair-stepped delta morphology (Muto & Steel 2001;Ritchie et al. 2004b;Villiers et al. 2013). Forced regression is defined as basinward shoreline retreat during relative base-level fall, whereas normal regression mayoccur during the base-level lowstand, rise and highstand, if the sediment supply exceeds the rate at which accommodation space is created (Posamentier & Morris 2000). Alluvialplain aggradation and delta-front progradation commonly accompany normal regression, whereas fluvial incision and sediment bypass occur during forced regression (Posamentier & Morris 2000; Ritchie et al. 2004a, b;Strong & Paola 2008;Catuneanu et al. 2011), leading to a rapid basinward stepping of delta lobes (Fig. 1). A possible genetic link between the delta-front morphodynamic responses to baselevel changes and the delta-slope sedimentation processes may help in the recognition of a hidden record of base-level change if the topset-foreset transition zone is eroded (Gobo et al. 2014(Gobo et al. , 2015.

Study area
The study area is located south of the North German Lowlands (Fig. 2). Luminescence data of ice-marginal deposits (Roskosch et al. 2015;Lang et al. 2018) point to several ice advances into this area during the Middle Pleistocene (Marine Isotope Stages MIS 12 to 6). The blocking of river valleys by the Middle Pleistocene Elsterian and Saalian ice sheets led to the repetitive formation of proglacial lakes (Eissmann 2002;Winsemann et al. 2007Winsemann et al. , 2009Roskosch et al. 2015;Lang et al. 2018). These proglacial lakes were characterized by overall water rises during ice advances, when lake-overspill channels were successively closed. Maximum lake levels of~200 m a.s.l. were reached during the Saalian glaciation, with lakelevel rises of up to 150 m within a few hundreds to thousand years Lang et al. 2018). During deglaciation, the lakes catastrophically drained due to the renewed opening of lake outlets, which caused rapid, high-magnitude lake-level falls in the range of 20-80 m within perhaps a few weeks (Meinsen et al. 2011;Winsemann et al. 2011Winsemann et al. , 2016Lang et al. 2018). The lake-level history of glacial lakes along the Elsterian ice-sheet margins (MIS 12 and 10) in northern central Europe is less well studied. The maximum lake levels in eastern Germany were probably similar to those of the Saalian glaciation, controlled by the topographic height of lake-overspill channels . The Elsterian lake levels of glacial Lake Leine probably reached 155 m a.s.l. (Roskosch et al. 2015), corresponding to a lake-level rises of approximately 80 m. Estimated lakelevel falls during deglaciation were in the range of 20-25 m. It is not known if larger glacial lakes existed in the Weser Valley during the Elsterian glaciations.
The ice-marginal delta systems are relatively small, ranging in size from~1.5 to~5 km 2 . Their thickness varies between~35 and~70 m. The Gilbert-type deltas are commonly located in front of steep mountain ridges and are fed by bedrock-feeder channels (Winsemann et al. 2007. In some cases, subaqueous ice-contact fans are downlapped, onlapped or overlain by Gilbert-type delta systems and/or shoal-water deltas (Winsemann et al. 2009). New outcrops reveal the presence of abundant bedforms deposited by supercritical density flows (Lang et al. 2017b). Gravelly shoal-water deltas formed during lake-level fall when water depths became low (Roskosch et al. 2015).

Geomorphology and sedimentology of delta systems
High-resolution digital elevation models (10-m grid, vertical resolution: AE0.5 m) were analysed in a geographical information system (ArcGIS). The geomorphology of the Freden delta has been reconstructed from old topographic maps (1901/1937) with ArcGIS software (3ArcGIS Version 10.3.1, Esri, Redlands, CA, USA).
Outcrop and borehole datawere studied to reconstruct the sedimentary facies and depositional architecture of ice-marginal delta systems. Vertical logs were measured at the scale of individual beds, noting grain size, bed thickness, bed contacts, bed geometry, internal sedimentary structures and palaeocurrent directions. Photomosaics of largeroutcropswere used for the interpretation of architectural elements. The terminology for gravel characteristics is after Walker (1975).

Ground-penetrating radar profiles
Ground-penetrating radar (GPR) was used to delineate architectural elements. These data provide a bridge between outcrop-facies architecture and the larger-scale   Winsemann et al. (2007Winsemann et al. ( , 2009Winsemann et al. ( , 2011 and Lang et al. (2018). The DEM is based on Copernicus data and information funded by the European Union (EU-DEM layers). Be = Betheln delta; Bo = Bornhausen delta; C = Coppenbr€ ugge subaqueous fan and delta complex; E = Emme delta; F = Freden delta; G = Großsteinberg delta; K = Karsdorf delta; M = Markendorf delta; P = Porta subaqueous fan and delta complex; W = W€ unsch delta; Z = Zeuchfeld delta. B-E. Geomorphology of the Porta subaqueous fan and delta complex, Emme delta, Betheln delta and Freden delta. Geomorphological profiles (I-I 0 ) are shown below each delta map. The DEMs of the Porta subaqueous fan and delta complex, Emme delta and Betheln delta are based on data from the Bezirksregierung K€ oln (10-m grid, vertical resolution: AE0.5 m) and LGN Hannover (10-m grid, vertical resolution: AE0.5 m). The DEM of the Freden delta is reconstructed from old topographic maps (1901/1937) with ArcGIS software. Contour lines are in 5-m intervals. [Colour figure can be viewed at www.boreas.dk] delta architecture mapped from shear-wave seismic profiles. The GPR device used was a GSSI SIR-3000 (Geophysical Survey Systems Inc. (GSSI), Nashua, NH, USA) with 200 and 400 MHz shielded antennas. Radar traces were collected every 5 cm along the profile and the data processing comprised dewowing, static correction, amplitude balancing by spherical divergence compensation and application of an exponential gain function, bandpass filtering and migration. The vertical resolution is~5-10 cm. The lateral resolution is 30-50 cm near the surface and~0.8-1.1 m at 5-m depth. The radar facies were defined on the basis of the external geometry and the internal reflector patterns (Gutsell et al. 2004;Neal 2004;Lee et al. 2007;Eilertsen et al. 2011).

Shear-wave seismic profiles
The larger-scale delta architecture was mapped from high-resolution shear-wave seismic profiles. The vertical resolution is up to~0.5 m. The lateral resolution is about 0.5 m near the surface and decreases to~12 m at 50-m depth. For all surveys presented here we combined a shear-wave land streamer with the micro-vibrator ELVIS (developed by LIAG) operating in transverse horizontal (SH) mode. Details on data acquisition and processing are given in Winsemann et al. (2011) and Roskosch et al. (2015).
The seismic facieswere defined on the basis of the external geometry and internal reflector patterns (Mitchum et al. 1977;Posamentier & Vail 1988). The seismic attributes amplitude and continuity were used for the analysis of reflector patterns (Bullimore et al. 2005). Geomorphology and large-scale depositional architecture of delta systems Four delta complexes were selected for this study, referred to as the Porta delta, Emme delta, Betheln delta and Freden delta ( Fig. 2A-E). These selected field cases are considered to be representative of the delta styles in glacial lake basins affected by rapid base-level change.

The Porta delta
The Saalian Porta subaqueous fan and delta complex is located at the northern margin of glacial Lake Weser. It has a bedrock-feeder channel, is approximately 6.2 km long and 5.3 km wide and has a radial lobate shape ( Fig. 2A, B). The delta system downlaps and onlaps the truncated subaqueous ice-contact fan subaqueous icecontact fan (Winsemann et al. 2009). On top of the truncated fan, a broad glacifluvial delta plain and shoalwater delta mouthbars developed, which fed the marginal Gilbert-type deltas ( Fig. 3A; Winsemann et al. 2009;Lang et al. 2017b). The Porta delta complex is up to~40 m thick and has a stair-stepped profile with two plains at~115 m and~95 m a.s.l. (Fig. 2B). It formed during an overall lake-level fall, punctuated by lower-magnitude lake-level fluctuations.

The Emme delta
The Saalian Emme delta is located at the northern margin of glacial Lake Weser. It has a deep bedrockfeeder channel, a radial, lobate shape and is about 2 km long, 1.8 km wide and up to 70 m thick ( Fig. 2A, C). Luminescence ages point to a deposition during MIS 6 . It overlies glaciolacustrine mud or Jurassic bedrock, forming a concave ramp, dipping at up to 13°. It has a stair-stepped profile with two plains at~128 m and~155 m a.s.l. The northeastern upper portion of the delta is characterized by a central, trumpet-shaped, up to 20-m-deep valley that rapidly shallows down-slope. The proximal valley has a sharp, steep western margin that can be traced for~500 m. In contrast, the eastern valley margin is less well developed and the valley opens rapidly towards the southeast. In front of this incised valley, depositional lobes with a telescoping morphology are developed. The margin of the Emme delta complex displays a radial pattern of ridges and smaller erosional valleys (Fig. 2C). The deposits of the Emme delta record one major transgressive-regressive cycle, punctuated by lower-magnitude lakelevel fluctuations . The oldest depositional units record the transgressive phase and are characterized by back-stepping delta bodies, decreasing upwards in grain size, thickness and lateral extent (Fig. 3B, seismic units 1-4). During a subsequent series of lake-level falls forced-regressive basinward stepping delta lobes formed that overlie and downlap the transgressive deposits (Fig. 3B, seismic units 5-9).

The Betheln delta
The Elsterian Betheln delta is located at the southwestern slope of the Hildesheimer Wald Mountains ( Fig. 2A, D) and has been deposited into glacial Lake Leine. Luminescence ages point to a deposition during MIS 12 (Roskosch et al. 2015). The fan-shaped, radial sediment body is about 3.5 km long, 1.5 km wide and up to 35 m thick. It overlies glaciolacustrine mud or Mesozoic bedrock, forming a ramp inclined at up to 6°. The feeder system consists of several parallel bedrock channels (Fig. 2D). The delta has a stair-stepped profile with three marked plains at~115,~135 and~155 m a.s.l. (Fig. 2D). The margin of the delta displays a radial pattern of ridges and smaller erosional valleys (Fig. 2D). The southwestern margin is eroded by the River Leine. The delta complex formed during two transgressive-regressive cycles (Roskosch et al. 2015). Deposits of the first transgressive-regressive cycle are characterized by a series of basinward stepping depositional lobes (Fig. 3C, seismic units 1-5), recording the lake-level highstand and subsequent lake-level fall. The second transgressiveregressive cycle is represented by seismic units 6-9. During rapid lake-level rise, the older delta units of the first transgressive-regressive cycle were onlapped and overlain by rapidly landward stepping delta lobes (seismic units 6 and 7). During subsequent high-magnitude lake-level falls basinward stepping delta lobes formed that downlap and overlie the older delta deposits (seismic units 8 and 9). The youngest delta units consist of shoalwater mouthbar deltas in the distal portion of the delta complex that downlap the steeply dipping Gilbert-type deltas.

The Freden delta
The Saalian Freden delta is located at the northern margin of glacial Lake Leine. It overlies a salt structure, the so-called Leine anticline (Winsemann et al. 2007;Brandes & Tanner 2012), and forms an isolated sediment ridge between two bedrock highs ( Fig. 2A, E) This sediment ridge is approximately 1 km wide, 1.5 km long and up to 60 m thick. Towards the northwest, the iceproximal slope has a concave-up profile inclined at 10°-2°; towards the southeast, the delta system has a multitongue-like lobate shape with a stair-stepped profile. Major plains occur at~180,~160 and~130 m a.s.l. (Fig. 2A, E). The depositional architecture points to the existence of two genetically different delta bodies, which are probably related to two transgressive-regressive cycles during MIS 6 and 8 (Roskosch et al. 2015). The older sand-rich delta deposits (MIS 8) were probably shed from the northeast via a bedrock feeder channel. The shear-wave seismic line (Fig. 4) shows a series of laterally and vertically stacked depositional lobes that formed during overall transgression. Within these delta deposits numerous shear-deformation bands are developed (Brandes & Tanner 2012;Brandes et al. 2018).
During the second ice advance (MIS 6), an ice-contact delta formed in front of an ice lobe that terminated in the lake. These delta deposits contain flow-till layers and glaciotectonic deformation structures, and downlap or unconformably overlie the older delta system (Roskosch et al. 2015).

Delta facies associations
The studied deltaic systems comprise Gilbert-type deltas and shoal-water mouthbar deltas. This section summarizes the sedimentary facies associations (FA) and geometrical Fig. 4. Shear-wave seismic profile of the Freden delta showing a series of laterally and vertically stacked deltas lobes, deposited during overall transgression and lake-level highstand. A. Uninterpreted seismic profile. B. Interpreted seismic profile. C. Velocity coded seismic profile. The shearwave interval velocity points to three major vertically stacked delta units that differ in velocity. Lower interval velocities and discontinuous, lowamplitude reflectors correlate with coarser-grained sand and pebbly sand. Higher interval velocities and higher amplitude, more continuous reflectors correlate with finer-grained, probably more compacted delta lobe deposits. [Colour figure can be viewed at www.boreas.dk] features of these delta systems. The delta deposits include a large variety of sedimentary facies (F), representing deposition from low-and high-energy tractional flows, debris fall, debrisflows, and sustained or surge-type supercritical to subcritical turbidity currents. Sediment clasts consist mainly of reworked fluvial material and poorly sorted, angular debris derived from the steep Mesozoic bedrock slopes. Clasts of a Scandinavian/Baltic provenance account for approximately 6-16%. All sedimentary facies (F) are briefly summarized in Table 1; facies associations (FA) are summarized in Table 2.

Gilbert-type deltas (FA1)
The Gilbert-type deltas are characterized by steeply dipping gravelly or sandy foresets that either pass tangentially into relatively flat-lying finer-grained bottomset facies or overlie bottomsets with an angular contact. In outcrops, the foreset tops are commonly bounded by erosional surfaces and no foreset-topset transition is recognizable. Erosional surfaces are related to the formation of long-wavelength bedforms, distributary channels or delta-top incised valleys. In seismic profiles, foreset-topset transitions display rising sigmoidal, smooth-topped subhorizontal or falling stepped-topped patterns (Winsemann et al. 2007(Winsemann et al. , 2009Roskosch et al. 2015).
In outcrop, the delta-plain deposits (FA1.1.1) have an overall sheet-like geometry and consist of 1-6 m thick trough cross-stratified sand, pebbly sand and gravel (facies St/Gt),which fill shallowlenticularchannels(upto 10 m wide and 2 m deep) with a nested offset stacking pattern. In seismic profiles, the delta-plain deposits are characterized by horizontal high-amplitude reflectors (Winsemannet al.2011). Inoutcrop,isolated metre-scale scours on top of truncated foresets are filled with gravelly backsets (facies Gbl).
Long-wavelength bedforms (FA1.1.2), deposited from supercritical tractional flows, consist of up to 10-m-high, slightly asymmetrical to symmetrical sediment waves with wavelengths of 60-90 m. In seismic profiles, the wave-like structures are characterized by internal convex-up parallel high-amplitude reflectors (Fig. 5A). Bedforms with shorter wavelengths of 38-45 m are associated with deep scours, filled with foresets ( Fig. 5B). Based on their wavelengths and asymmetry these deposits may represent either large antidunes or net-depositional cyclic steps (Kostic 2011;Winsemann et al. 2011;Muto et al. 2012;Cartigny et al. 2014). The downflow alternation with irregularly spaced scours (Fig. 5B) indicates cyclic steps or chutes-and-pools with superimposed antidunes Zhong et al. 2015;Lang et al. 2017a). Scours are filled laterally with foresets, which more commonly occur in chutes-andpools (Lang & Winsemann 2013;Cartigny et al. 2014). However, cross-strata backsets are more characteristic of cyclic-step bedforms and the latter cannot thus be ruled out. Net-erosional cyclic steps would produce trains of scours, which might have been filled subsequently by lower-energy currents during the final stage of flow. Isolated scour fills with backsets are interpreted as the preserved hydraulic jump zone of cyclic steps (Muto et al. 2012;Cartigny et al. 2014) that probably formed on the delta plain during major meltwater discharges.
Distributary-channel fills (FA1.1.3) are lens-shaped in cross-section and appear as vertically and laterally stacked lenticular deposits that are often organized into larger scale channel complexes, up to 8 m thick and~100 m wide. Individual channels are up to 70 m wide, 1.5-5 m deep and have aspect (width/depth) ratios >14:1. In shearwave seismic profiles, larger distributary channels are recognizable as concave-up, lenticular features with highamplitude internal reflectors . At the base of these channels gravel lags are common, composed of clast-supported cobble to boulder gravel (facies B). These channel-floor deposits are overlain by trough cross-stratified cobble to pebble gravel, pebbly sand and sand (facies Gt/St). Towards the channel margins, cross-sets are finer-grained and thinner and often form climbing cosets (facies Scd). Climbing dunes formed under high suspension fall-out rates  and may indicate the channel-mouth zone (Ghienne et al. 2010;Carvalho & Vesely 2017). The upper channel fills commonly consist of fine-grained planarparallel stratified and climbing-ripple cross-laminated fine-grained sand, silt and mud (facies Sr/Fl), indicating channel abandonment .
Incised-valley fills (FA1.1.4) are approximately 25 m deep and 25-150 m wide, commonly flat-based and cut deeply into delta topset and foreset deposits. They form large-scale U-shaped isolated features with aspect (width/depth) ratios of 1:1 to 6:1. In seismic profiles, incised valleys are concave-up erosional features with transparent internal reflectors ). The axial valley-fill deposits consist mainly of thickbedded planar and trough cross-stratified sand, pebbly sand and gravel (facies St/Gp/Gt), deposited from turbulent subcritical tractional flows. Low-angle crossstratified or sinusoidally stratified sand, pebbly sand and gravel beds (facies Sl/Gl) indicate deposition from antidunes during supercritical flow conditions. As in the delta-plain distributary channels cross-sets are finergrained and thinner towards the channel margins and often form climbing cosets (facies Scd/Sr). The internal Massive or inversely graded matrix-or clast-supported pebble to boulder gravel. The matrix is fine-to coarse-grained sand. Long axes of large outsize clasts may be orientated parallel to dip and may show a steep upslope dipping a(p) a(i) fabric. Bed contacts are sharp and non-erosional Deposition from non-cohesive debrisflows. The steep-clast fabric indicates laminar shear during or immediately after the flow's stop Gp/Gt Clast-supported pebble to cobble gravel with planar or trough cross-stratification. The matrix is fine-to medium-grained sand. Bed contacts are erosive Deposits of migrating 2D and 3D dunes or transverse bars. The deposition occurred from tractional bedload flows in channels on the delta plain or sustained turbidity currents on the delta slope Gl Low-angle cross-stratified gravel. The matrix is medium-to coarse-grained sand Deposits of breaking antidunes. Deposition from supercritical tractional flows in channels on the delta plain or supercritical turbidity currents on the delta slope Gbl Backset cross-stratified gravel that laterally may pass into lowangle cross-stratified or sinusoidally stratified sand. Backsets occur in regularly spaced scours or over the entire bed length. The gravel typically shows an upslope dipping steep-clast fabric (a(p) a Deposits of cyclic steps. Deposition from surge-type or sustained supercritical turbidity currents on the delta slope

Smg
Pebbly fine-to coarse-grained massive, inversely or normally graded sand. Clasts are commonly pebble-to cobble-sized. Bed contacts are sharp or erosive Deposition from sandy debrisflows by freezing or turbidity currents on the delta slope Sbl Backset cross-stratified pebbly sand and sand that laterally may pass into low-angle cross-stratified or sinusoidally stratified sand. Backsets occur in regularly spaced scours or over the entire bed length. Some sandy scour fills are massive, diffusely graded or deformed by dewatering structures and may pass upslope into backset cross-stratification Deposits of cyclic steps. Deposition occurred from supercritical turbidity currents on the delta slope. Scour fills with deformed strata and dewatering structures reflect the hydraulic-jump zone of cyclic steps where rapid suspension fall-out and liquefaction of deposits occur Sl Low-angle or sinusoidally cross-stratified pebbly sand and sand. Beds may fine or coarse upwards and have internal truncation surfaces. At the base small scours filled with pebbles may occur. Upsetion the thickness of bedsets commonly increases and sigmoidal bedforms are preserved. Bed contacts are sharp or gradational Deposits of breaking and stationary antidunes. Deposition from supercritical tractional bedload flows in channels on the delta plain or supercritical turbidity currents on the delta slope Ssi Sigmoidally cross-stratified sand and pebbly sand with welldeveloped topset, foreset and bottomset geometries. Bed contacts are sharp or erosive Deposits of migrating humpback dunes. Deposition from transcritical turbidity currents on the delta slope Sp/St Fine-to coarse-grained sand and pebbly sand with planar or trough cross-stratification. Beds partly form climbing bedsets. Bed contacts are sharp or erosive Deposits of 2D and 3D dunes or transverse bars. Deposition occurred from subcritical tractional bedload flows in channels on the delta plain or sustained turbidity currents on the delta slope Scd Planar, trough or sigmoidally cross-stratified pebbly sand and sand forming climbing bedsets. Foreset beds are partly oversteepened and contorted and may pass updip into low-angle cross-stratified sand (Sl). Bed contacts are erosive to gradational Deposits of migrating 2D and 3D dunes or humpback dunes. Deposition from subcritical to transcritical tractional flows in channels on the delta plain or subcritical to transcritical turbidity currents on the delta slope. Climbing bedforms indicate high suspension fall-out rates, partly under hydraulic-jump conditions Sr Fine-to coarse-grained (climbing) ripple cross-laminated sand.
Beds are planar or trough cross-laminated and commonly show a fining-upward where a lamination with eroded ripple stoss sides passes upwards into lamination with preserved stoss sides and into draping lamination. Bed contacts are sharp, erosive or gradational Deposits of 2D and 3D ripples. Deposition from subcritical tractional flows in channels or interchannel areas on the delta plain or sustained subcritical turbidity currents on the delta slope and prodelta. Climbing bedforms indicate high suspension fallout rates

Fl
Normally graded or massive sand that fines upwards into planarparallel laminated and ripple-cross laminated medium-to finegrained sand and silt, planar-parallel laminated or massive silt, mud or clay. Beds are most commonly 'incomplete' and contain both 'top-absent' or 'base-absent' successions. Bed contacts are sharp or erosive Deposition from subcritical tractional flows or suspension fallout on the delta plain or from waning surge-like subcritical turbidity currents on the delta slope or prodelta Table 2. Facies associations of (A) Gilbert-type delta and (B) shoal-water delta deposits.

Delta topsets
Delta-plain deposits (FA1.1.1) Trough cross-stratified sand, pebbly sand and gravel (facies St, Gt). Troughs are 0.3-1.5 m wide, 0.15-0.5 m thick and fill shallow lenticular channels Deposition by migrating 3D dunes in a lower flow regime of uni-directional currents (Harms et al. 1975;Ghienne et al. 2010;Winsemann et al. 2011) Long-wavelength bedforms (FA1.1.2) Long-wavelength bedforms consist of up to 10 m high, slightly asymmetrical to symmetrical sediment waves with wavelengths of 60-90 m. The asymmetric bedforms have slightly steeper upflow (stoss) slopes than the downflow ( Deposition by migrating 3D dunes in lower flow regime uni-directional currents (Harms et al. 1975;Ghienne et al. 2010) and by antidunes during supercritical flow conditions (Fielding 2006). Thick planar cross-stratified gravel beds resulted from the migration of gravel bars (Massari & Parea 1990;Browne & Naish 2003). Climbing dunes at the valley margins indicate lower flow velocities and high suspension fallout rates  Delta foresets Foreset-bed packages I (FA1.2.1) Massive or inversely graded matrix-or clast-supported pebble to boulder gravel with sharp nonerosive basal contacts (facies Gmg). The matrix is medium-to coarse-grained sand. Many beds show upslope-dipping internal shear planes and an imbricated a(p) a(i) fabric. Bed thickness is 10-60 cm. Beds commonly have sharp, mostly non-erosional contacts. Erosion surfaces are commonly draped by thin-to medium-bedded (2-20 cm) normallygraded sand (facies Smg, Fl) or low-angle cross-stratified pebbly sand beds (facies Sl) that laterally grade into fine-grained bottomset beds Deposition from cohesionless debrisflows (Nemec 1990;Sohn et al. 1997). The thin sand or pebbly sand beds that drape major erosional surfaces indicate deposition from subcritical and supercritical turbidity currents (Kostic et al. 2002;Winsemann et al. 2011) Foreset-bed packages II (FA1.2.2) Medium-to thick-bedded (10-60 cm) massive, normally or inversely graded matrix-or clastsupported pebble to boulder gravel (facies Gmg) and pebbly sand (facies Smg). The sorting is generally poor and the matrix consists of medium-to coarse-grained sand. Long axes ofoutsized clasts are often aligned parallel to the dip of bedding planes. Beds contacts are sharp, uneven and mostly non-erosional. Some pebbly sand and gravel beds are low-angle cross-stratified (facies Gl, Sl), 10-20 cm thick, and show erosional basal contacts. Occasionally isolated scours, 0.1-0.3 m deep and 0.5-1 m wide, filled with sandy or gravelly backsets occur (facies Sbl, Gbl).
Clast-supported open-work boulder to pebbly gravel (facies Go) in the lower foreset or toeset area are 0.6-3 m long in dip direction and 0.05-1.5 m thick. Theyoften display avertical normal grading and a lateral grading of a coarse head down-dip into an upslope fining tail Deposition from non-cohesive debrisflows, sandy debrisflows (Sohn et al. 1997;Nemec et al. 1999), debris fall (Nemec 1990;Sohn et al. 1997; Uli cn y 2001) and surge-type supercritical turbidity currents (Lang et al. 2017b) Foreset-bed packages III (FA1.2.3-FA1.2.5) FA1.2.3: Backset cross-stratified gravel and pebbly sand, alternating with low-angle crossstratified sand and pebbly sand, climbing-ripple cross-laminated or massive sand and silt, forming small-scale (0.9-1.5 m) fining-upward successions. At the base, erosive-based gravel and pebbly sand beds frequently show backsets that fill scours with scooped basal erosion surfaces, 0.9 m to >4 m wide and 0.1-0.7 m deep (facies Gbl). The gravel commonly shows a steep upslope dipping a(p) a(i) or a(t) b(i) fabric. Some sandy scour fills are massive, diffusely graded or deformed by dewatering structures (convolute bedding and clastic dykes) and may pass upslope into backset cross-stratification. Laterally scour fills may pass into sheet-like low-angle crossstratified or sinusoidal bedforms (facies Sbl). The overlying low-angle cross-stratified sand beds Tractional deposition from waning, surge-type supercritical to subcritical turbidity currents that produce small-scale fining-upward successions Ventra et al. 2015;Lang et al. 2017a, b) (continued) Interpretation are 5-15 cm thick, have erosive bases and may be draped by thin (0.5-1 cm) massive silty finegrained sand layers. They have internal truncation surfaces and may show small concave-up scours (5-10 cm wide and 2-3 cm deep) at the base filled with pebbles (facies Sl). Upsetion the thickness of bedsets commonly increases and sigmoidal bedforms are preserved. Occasionally gravelly or sandy sigmoidally cross-stratified deposits occur (facies Ssi). The uppermost portion of the fining-upward successions may consist of thin beds (1-10 cm thick) with climbing ripple cross-laminated silty sand or massive silt and mud (facies Fl). The small-scale fining-upward sequences may be organized into larger-scale (2-3 m) coarsening or fining-upward cycles, characterized by the thickness and abundance of gravel beds FA1.2.4: Backset cross-stratified sand and pebbly sand, alternating with low-angle crossstratified, sigmoidally cross-stratified and planar and trough cross-stratified sand and pebbly sand. Pebbly sand beds frequently show backsets that fill scourswith concave-up geometries, 0.6 -6 m long and 0.06-0.7 m deep (facies Sbl). Individual beds are 0.2-1.6 m thick, commonly fine upwards and occur over the entire foreset length. Laterally and vertically backset beds may pass into sheet-like low-angle cross-stratified (facies Sl) or sinusoidal bedforms (facies Ssi), 0.2-0.7 m thick, forming dm-scale fining-upward successions. Perpendicular and oblique to flow these deposits appear as shallow troughs, filled with concentric to low-angle cross-stratified pebbly sand and sand. Finer-grained sandy beds commonly display dune-scale planar and trough cross-stratification (facies Sp, St). On average beds are 0.1-0.7 m thick. Bed contacts are sharp erosional. In the delta-toe zone well-preserved deposits of sigmoidal humpback dunes are often developed, which show typical tripartite geometries with topsets, foresets and bottomsets (facies Ssi) Tractional deposition from sustained supercritical to subcritical turbidity currents Lang et al. 2017a, b). The formation of sigmoidal humpback dunes requires highly aggradational conditions (Fielding 2006;Lang & Winsemann 2013;Cartigny et al. 2014), which prevail in the deltafoot zone FA1.2.5: Climbing-ripple cross-laminated sand with intercalations of lenticular massive or backset cross-stratified pebbly sand and sand beds. Beds mainly consist of medium-to very thick-bedded (0.1-1.8 m) fine-to medium-grained climbing-ripple cross-laminated sand (facies Sr). Some beds contain scattered pebbles. Beds often show a fining-upward where a lamination with eroded ripple stoss sides passes upwards into lamination with preserved stoss sides and into draping lamination and very thin-bedded mud and clay beds. Ripples either migrate down-slope or upslope. More rarely, thin-to medium-bedded sand, silt and mud alternations with Bouma Ta-d divisions (facies Fl), large-scale cross-stratified pebbly sand beds (facies Sp,St) or lenticular intercalations, 1.2-10 m wide and 0.2-0.8 m thick, occur that consist of massive, diffusely graded, deformed or backset cross-stratified pebbly sand and sand beds (facies Sbl) Deposition by sustained subcritical turbidity flows, which produce beds without significant vertical variation in grain size (Kneller & Branney 1995;Winsemann et al. 2007). During higher flow conditions dune-scale crossstratification and cyclic steps with backset cross-stratification accumulated on the lower delta slope Lang et al. 2017b). Upslope migrating ripples may indicate the zone of flow transition of jets emerging from delta-plain channels (Jopling 1965) Chute fills: In foreset packages III (FA1.2.3 and FA1.2.4) lenticular chute fills are common. The fill consists of trough cross-stratified gravel (facies Gt), overlain by massive to diffusely stratified sand (facies Smg/Sl) or cross-stratified and/or ripple cross-laminated sand (facies Sp/St/Sr). Troughs of the gravelly bedforms are 1-4 m wide and 0.2-2.5 m deep and may contain cobblesized intraclasts. Grain size and matrix content vary between individual troughs. Some coarsergrained trough fills display open framework. Occasionally found are gravelly scour fills with backset cross-stratification (facies Gbl). The laterally more persistent fine-to coarse-grained sand beds are 0.4-0.7 m thick and fine upwards. At the base of the chutes often a boulder to cobble gravel lag (facies B) occurs. Towards the top and channel margin the thickness of crosssets commonly decreases (0.2-0.6 m) and climbing cosets are often developed. Foreset beds of the climbing (humpback) dunes (facies Scd) are partly oversteepened and contorted and pass updip into low-angle cross-stratified sand (facies Sl) Tractional deposition from supercritical and subcritical turbidity currents (Mulder & Alexander 2001;Winsemann et al. 2009). High sedimentation rates under hydraulic-jump conditions are indicated by the formation of climbing humpback dunes with oversteepened and contorted foresets Lang & Winsemann 2013). The oversteepened and contorted dune foreset beds indicate liquefaction-induced slope collapse processes caused by rapid loading (Owen 1996) Delta bottomsets FA1.3.1 Low-angle cross-stratified medium-to coarse-grained sand and pebbly sand (facies Sl), interbedded with massive or inversely graded pebble to cobble gravel with non-erosive bases (facies Gmg). Bed thickness ranges between 0.1 and 0.3 m. Isolated larger clasts or small gravel clusters occur in distinct sand and pebbly beds. The long axes are commonly orientated parallel to the dip of bedding planes. The low-angle cross-stratified pebbly sand and sand beds may pass downflow into climbing-dune cross-stratification (facies Scd), forming 0.2-0.3 m thick cosets. Occasionally small isolated scours occur (0.5-1 m long and 0.1-0.15 m deep) that are filledwith sandy backsets or foresets (facies Sbl). In finer-grained, more sand-rich bottomset deposits lowangle and sigmoidally cross-stratified sand and pebbly sand (facies Sl and Ssi) may alternate with thin-to medium bedded (0.1-0.3 m) climbing-ripple cross-laminated sand and silt (facies Sr) Deposition from cohesionless debrisflows, debris fall and supercritical to subcritical turbidity currents, partly triggered by the release of limited sediment volumes by discrete failures of upper delta-slope deposits (Nemec 1990;Sohn et al. 1997;Nemec et al. 1999;Winsemann et al. 2011;Gobo et al. 2014).
Hydraulic-jump conditions in the prodelta zone led to the formation of isolated scours and the deposition of small-scale climbing dunes (Nemec et al. 1999;Winsemann et al. 2007Winsemann et al. , 2011 FA1.3.2 Thin-to thick-bedded (0.1-0.5 m) climbing humpback-dune assemblages (facies Scd) passing downflow into thin-to medium-bedded (0.05-0.3 m) climbing-ripple trough cross-laminated fine-to coarse-grained sand (facies Sr) and intovery thin-to thin-bedded (2-10 cm) alternations Deposition by turbidity currents under hydraulic-jump conditions during flow expansion at the mouth of a channel or slope break (Nemec et al. 1999;Macdonald et al. (continued) architecture of the incised-valley fills is commonly characterized by an amalgamated vertical stacking of channel-fill deposits in the valley axis, and an onlapping, laterally offset stacking at the valley margins, along which high-angle slide scars and large slide blocks can be found (Winsemann et al. 2007. These intra-valley channels are~5 to >40 m wide,~2-10 m deep and show a fining-upwards-trend with a basal gravel lag (facies B), overlain by a succession of cross-stratified gravel and pebbly sand (facies St/Gp/Gt), ripple-cross-laminated sand (Sr) and thinly interlayered silt and mud (facies Fl). The multistorey and heterogeneous infills of these intra-valley channels indicate several phases of channel scouring and deposition, related to channel migration and/or variations in meltwater discharge (Olariu & Bhattacharya 2006;Winsemann et al. 2007Winsemann et al. , 2011. Towards the top of an incised-valley fill the channelmargin deposits locally pass into thinner-bedded, sheetlike deposits (facies Fl), onlapping directly the truncated foreset. These overbank deposits reflect an increased range of the lateral shifting of glacifluvial channels with the decreasing valley accommodation.
Foreset facies association (FA1.2). -Delta foresets have thicknesses between 5 and 25 m and foreset beds are inclined between 5°and 34°. In strike sections, the foreset deposits form laterally and vertically stacked mounds, 15-360 m wide. The sedimentary facies of the delta foresets include a wide range of gravity-flow deposits that tend to form three distinct facies assemblages: (I) foreset-bed packages dominated by debrisflow deposits (FA1.2.1); (II) foreset-bed packages dominated by debrisflow and debris-fall deposits (FA1.2.2); and (III) foreset-bed packages deposited by supercritical and subcritical low-and high-density turbidity currents (FA1.2.3-FA1.2.5). These foreset-bed packages are often separated from one another by erosional surfaces that dip less steeply and differ in dip directions. Chute channels, 8-60 m wide and 1-5 m deep, are common in foreset packages II and III and are mainly filled with deposits of low-and high-density turbidity currents (Figs 6-8).
Foreset-bed packages I, dominated by debrisflow deposits (FA1.2.1), consist of massive or inversely graded, matrix-or clast-supported, pebble to boulder Interpretation of massive or planar parallel-laminated clay and planar-parallel or climbing-ripple crosslaminated silt and fine-to medium-grained sand (facies Fl). Cm-scale convolute bedding, ball and pillow structures, and flame structures are common. Bed contacts are erosional to gradational 2009; Winsemann et al. 2011). The lateral facies transition from climbing dunes into finergrained facies Sr and Fl records waning flow conditions and deposition from diluted turbidity currents FA1.3.3 Thin-to medium-bedded (0.05-0.3 m) fine-to coarse-grained climbing-ripple cross-laminated sand (facies Sr). Some beds show a thin basal unit with planar-parallel lamination. Ripples are planar or trough cross-laminated and beds may show a fining-upward where lamination with eroded ripple stoss sides passes upwards into lamination with preserved stoss sides and into draping lamination. These beds may be intercalated with normally graded or massive sand that fines upwards into planar-parallel laminated and ripple cross-laminated medium-to finegrained sand and silt, laminated silt, and finally into laminated or massive mud or clay (facies Fl) Deposition from subcritical surge-type and sustained turbidity currents (Ashley et al. 1991;Kneller & Branney 1995;Mulder & Alexander 2001) FA2.1 Vertically stacked sets of flat-based, convex-up, planar to sigmoidally cross-stratified mediumto coarse-grained sand and pebbly sand (facies Sp). The overall grain size of the foresets decreases upwards. Foreset beds are commonly laterally graded and have dip angles of 5-30°. Upflow foreset beds pass into subhorizontally stratified sand. Downflow foreset beds prograde over a thin (1-2 cm) subhorizontal bottomset layer. The bedsets are 1.5-2 m thick and clinoforms are partly incised by small channels (2-3 m wide and 0.2-0.5 m deep) filled with trough cross-stratified medium-grained sand (facies St) Bedload deposition from inertia-dominated subcritical jets (Wright 1977;Postma 1990). The overall convex-up geometries, the downstream migration and the absence of major channels point to a distal delta mouthbar environment (Fielding et al. 2005;Lee et al. 2007) FA2.2 0.3-1.5 m thick trough-cross stratified cobble to pebble gravel (facies Gt), alternating with 0.1-0.7 m thick sigmoidally cross-stratified (facies Ssi), trough cross-stratified (facies St), planar cross-stratified (facies Sp), low-angle cross-stratified (facies Sl) and ripple cross-laminated (facies Sr) coarse-to fine-grained sand. The ripples partly form climbing bedsets. Larger-scale convex-up sigmoidally cross-stratified sandy bedforms, 1-2 m high, may pass downflow into climbing-ripple cross-laminated sand and onlap and drape coarser-grained convex-up bedforms, partly showing upstream accretion. These deposits are arranged into metre-scale (1.5 -2.5 m) fining-upward or coarsening-upward successions, bounded by major subhorizontal or slightly concave-up erosional surfaces. The bounding surfaces may be draped by thin layers of silty sand and partly show steep-flanked V-shaped scours (up to 0.5 m deep) at the base that are laterally filled with gravel. The overlying deposits may include large sandy intraclasts, up to 0.7 m in diameter. Flow directions are highly variable and show a dispersion of up to 90°B edload deposition from inertia-dominated supercritical to subcritical jets (Wright 1977;Postma 1990). The formation of V-shaped scours and intraclasts may be related to the formation of cyclic steps during supercritical flow conditions . Laterally overlapping delta lobes and a high dispersion of flow directions point to a proximal delta-front environment (Olariu & Bhattacharya 2006;Lee et al. 2007;Fidolini & Ghinassi 2016) gravel with sharp non-erosive basal contacts (facies Gmg; Fig. 6A). Many beds show upslope-dipping internal shear planes (Fig. 6B) and an imbricated a(p) a(i) fabric. In dip section these foreset beds are extensive, fairly tabular and overlie bottomset beds (FA1.3.2) with an angular contact (Fig. 6C). Internal erosion surfaces are scarce and typically have lower dip angles than the foreset bedding. Erosion surfaces are commonly draped by thin-to medium-bedded, normally graded sand (facies Smg/Fl) or low-angle cross-stratified pebbly sand beds (facies Sl), deposited from subcritical and supercritical turbidity currents. These beds laterally grade into fine-grained bottomset beds (facies Fl, FA1.3.2). In seismic profiles, foreset-bed packages I are commonly characterized by a hummocky, transparent reflector pattern .
In foreset-bed packages II, dominated by debrisflow and debris-fall deposits (FA1.2.2), grain size varies from sand, pebbly sand to gravel (Fig. 6E-G) and the sorting is generally poor. The foreset facies includes massive, normally or inversely graded, matrix-or clast-supported pebble to boulder gravel (facies Gmg) and pebbly sand (facies Smg). These beds commonly have sharp, uneven and mostly non-erosional contacts, which might follow the irregular surfaces of underlying coarse gravel beds. Some pebbly sand and gravel beds are low-angle crossstratified (facies Gl/Sl) and show erosional basal contacts. Occasionally found are isolated scours filled with sandy or gravelly backsets (facies Sbl/Gbl). Along dip, the beds are either laterally fairly persistent or pinching out within a few metres. The most characteristic features of these foreset packages are clast-supported open-work gravel lenses (facies Go) in the lower foreset or toeset area (Fig. 6G). Foreset-bed packages, dominated by such gravel lenses overlie sandy bottomset beds with an angular contact, whereas the sand-richer, less steeply dipping foreset beds locally pass down-slope into sandy or gravelly bottomset deposits (Fig. 6E-H; FA1.3.1). In seismic profiles, foreset-bed packages II are characterized by mainly discontinuous low amplitude reflectors Roskosch et al. 2015).
Foreset-bed packages III, deposited by supercritical and subcritical turbidity currents, comprise three different facies associations (FA1.2.3-FA1.2.5) that are characterized mainly by tractional bedforms. Common is a lateral fining from coarser-grained foreset-bed packages (FA1.2.3) to finer-grained foreset-bed packages (FA1.2.5). Foreset beds deposited by supercritical turbidity currents consist of laterally and vertically stacked cyclic step and antidune deposits (facies Gbl, Sbl, Sl). Metre-scale fining-upward sequences and the frequent intercalation of silt and mud drapes in FA1.2.3 indicate waning, surge-type turbidity currents (Fig. 6I-M). In contrast, FA1.2.4 was deposited by more sustained supercritical to subcritical turbidity currents, indicated by thick backsets and dune-scale foresets that occur over the entire delta foreset length. Finer-grained silt or mud drapes are absent in this facies association (Fig. 7C-F). In GPR profiles backsets are characterized by lenticular elements, which are characterized by  sigmoidal upslope reflectors with low to medium amplitudes (Fig. 8A, B). Farther basinwards foresetbed packages display a higher continuity, the reflector spacing decreases and scours with backsets become rare. In the delta-toe zone well-preserved deposits of sigmoidal humpback dunes are often developed (facies Ssi), indicating less powerful transcritical sustained turbidity currents and highly aggradational conditions (Fielding 2006;Lang & Winsemann 2013;Cartigny et al. 2014). The upward development from trough cross-stratified pebbly sand to preserved bedforms of finer-grained humpback dunes and antidunes may indicate flow thinning over aggrading beds leading to temporarily accelerating transcritical to supercritical flow conditions  grained intercalations of pebbly sand beds with backsets appear more transparent (Fig. 8C).
Laterally extensive, convex-up, sigmoidally crossstratified sand and pebbly sand (I) (facies Sp; FA2.1; Fig. 9A, B) form vertically stacked sets of large, flatbased mouthbars (FA2.1). Up-flow foreset beds pass into subhorizontally stratified sand. Downflow foreset beds prograde over a thin (1-2 cm) subhorizontal bottomset layer. Clinoforms are partly incised by small lenticular channels, filled with trough cross-stratified medium-grained sand (facies St). In GPR profiles mouthbars are characterized by basinward-dipping moderate-to high-amplitude, continuous reflectors in flow direction (Fig. 9A) and mounded bidirectionally downlapping reflections perpendicular or oblique to flow. Mounds are more than 25 m wide and in dip direction clinoforms can be laterally traced for more than 70 m (Winsemann et al. 2009;Lang et al. 2017b). Locally, sigmoidal geometries with transitions into bottomsets and topsets occur. The foresets are bounded by high-amplitude, gently landward dipping reflectors. Locally, small troughs and truncation of foreset boundaries can be observed. Channelized features are rare. They consist of lenticular elements, 10 m wide and up to 1 m deep, infilled by nested stacks of concentric or tangential reflectors. Up to four packages of mouthbar deposits are vertically stacked.
Facies association FA2.2 (II) is characterized by medium-to thick-bedded poorly sorted trough-and planar cross-stratified cobble to pebble gravel and sand (facies Gt, Ssi, St, Sp,), sigmoidally and low-angle crossstratified sand (facies Ssi, Sl) and ripple cross-laminated sand (facies Sr). Larger-scale convex-up bar elements partly show upstream accretion. These deposits are arranged into metre-scale fining-upward or coarseningupward successions, bounded by major subhorizontal or slightly concave-up erosional surfaces (Fig. 10A-E). The bounding surfaces may be draped by thin layers of silty sand (facies Fl) and partly show steep-flanked Vshaped scours at the base that are laterally filled with gravel. The overlying deposits may include large sandy intraclasts, up to 0.7 m in diameter (Fig. 10D, E). Flow directions are highly variable and show a dispersion of up to 90°. The overall geometry mapped from outcrops and GPR profiles perpendicular and oblique to flow is characterized by laterally and vertically stacked mounds, 6-25 m wide and 0.3-3 m thick (Fig. 9C). In flow direction deposits are characterized by gently basinward dipping clinoforms (Fig. 9D). The lower boundaries are concordant or downlapping. Internally, parallel, continuous reflectors dominate. Amplitudes are high to low. In general, smaller mounds are associated with higher amplitude reflectors. Concave-up channelized features are 3-8 m wide and 0.5-1 m deep. Internally they comprise laterally and vertically stacked lenticular elements with inclined-tangential reflectors. The mounded delta lobes are commonly top-preserved with mounded bar crests. Erosion is limited to the small-scale channel elements, which are incised into the upper parts of the delta lobes. Two to three lobe elements are vertically stacked. Compensational stacking is only observed for smallerscale lobe elements, occurring perched in the troughs between the larger-scale elements (Fig. 9C).
Laterally overlapping coarse-grained, poorly sorted shoal-water delta lobe deposits and a high dispersion of flow directions point to frequent autocyclic lobe switching and channel avulsion in a proximal delta-front environment (Olariu & Bhattacharya 2006;Lee et al. 2007;Fidolini & Ghinassi 2016). Initial mouthbars formed close to the channel axis, leading to flow splitting and the formation of new terminal distributary channels at different scales. Erosional surfaces with steep-flanked V-shaped scours and large intraclasts (Fig. 10D, E) are interpreted as bases of distributary channels (Olariu & Bhattacharya 2006;Lee et al. 2007). The formation of V-shaped scours and intraclasts may be related to the formation of cyclic steps during supercritical flow conditions (Strong & Paola 2008;Muto et al. 2012;. Fining-upward cycles may represent successive waning flows of major discharge events (Fielding et al. 2005). The slightly basinward dipping laterally extensive bounding surfaces (Fig. 9D) represent delta-lobe boundaries separating more distal from more proximal mouthbar lobes (Lee et al. 2007).

Deformation structures
The deformation structures within the ice-marginal deltas comprise both contractional and extensional features. Contractional structures are generally sparse in the studied delta systems. The faults have planar to slightly listric geometries with offsets in the range of metres to tens of metres. These thrusts commonly sole out into a basal detachment, which is controlled by lithological contrasts. The most characteristic deformation structures in the studied ice-marginal deltas are normal faults and shear-deformation bands. The normal faults commonly show synsedimentary activity and two different types of fault systems can be distinguished.
The first type are normal fault systems, which are restricted to the delta body. These faults have a slightly listric geometry and form small graben and half-graben systems (70-100 m wide), which locally show roll-over structures. Vertical offsets range between 2 and 15 m (Figs 3B, 4B). The fill of the half-grabens has a wedge-shaped geometry, with the greatest sediment thickness close to the fault, indicating synsedimentary activity. It is not clear if the faults sole out in a subhorizontal detachment . In outcrops bed displacements along normal faults are only in the range of a few centimetres to decimetres. Some delta-slope channel fills are bounded by high-angle (65-90°) gravitational synsedimentary normal faults with vertical offsets of 0.1-1.2 m (Winsemann et al. 2009). Fault systems, which are restricted to the delta bodies, are related to gravitational deformation, where extension in the upper parts of the delta body is compensated by contraction at the delta-toe (Bilotti & Shaw 2005;Bini et al. 2007;Brandes et al. 2007aBrandes et al. , 2011. The second type are normal fault systems that originated in the underlying Mesozoic bedrock and propagated into the overlying Pleistocene delta bodies. These faults are closely spaced (10-40 m) with vertical offsets of 2-15 m and may form small-scale graben or half-graben structures (120 m wide). Wedge-shaped geometries of halfgraben fills and thickening of reflector packages above the graben fills indicate synsedimentary activity (Figs 3B, 4B). Shear-deformation bands are 0-8 cm thick (mean 1.5 cm) and may displace beds by decimetres (Brandes & Tanner 2012). They form dense arrays of regularly spaced structures (Fig. 8A, B).
Fault systems that originated in the underlying Mesozoic bedrock and propagated into the overlying Pleistocene delta bodies indicate a Pleistocene reactivation of Mesozoic fault systems. This reactivation is probably related to the extension in the forebulge area of the advancing ice sheet, in combination with water and sediment loading. The dense arrays of shear-deformation bands are either related to salt movements and enhanced crestal collapse, or to reactivation of basement faults due to ice loading (Brandes & Tanner 2012;Brandes et al. 2018).
The frequent occurrence of bedforms deposited by supercritical turbidity flows along the foresets may be a characteristic feature of high-energy coarse-grained deltas (Massari 1996;Ventra et al. 2015;Dietrich et al. 2016;Lang et al. 2017b;Massari 2017). Commonly, isolated scour fills with massive, deformed or backset cross-stratified deposits were reported from the lower delta slope and delta-foot zone and related to hydraulic jumps at a break in the slope gradient, leading to rapid cut-and-fill processes (Clemmensen & Houmark-Nielsen 1981;Nemec et al. 1999;Winsemann et al. 2007;Gobo et al. 2014). However, bedforms of supercritical density flows may have been overlooked in the past and interpreted as deposits of cohesionless (sandy) debrisflows or suspension-fall-out deposits from sustained turbidity flows, because only recently have numerical simulations (Kostic 2011) and flume experiments ) allowed for a better understanding and recognition of these large-scale bedforms.

Base-level control on sedimentary facies, facies associations and stratal geometries
The sedimentary facies and facies associations defined from outcrop analysis were correlated with seismic units and subsequently assigned to base level (Fig. 11). The exposed delta sediments mainly comprise highstand, forced regressive and lowstand deposits, which record the phase of maximum lake level and subsequent lake drainage. Their features are considered as representative of delta styles in glacial lake basins, which are affected by rapid lake-level change. The good preservation of delta sediments during overall lake-level fall is related to the deglaciation stage, during which meltwater volumes, subglacial lake-outburst floods and sediment supply may increase and transfer large volume of sediments via highmagnitude discharges into deltaic environments (Evans & Clague 1994;Marren 2005;Ghienne et al. 2010).
Deposits of the overall transgressive phase during lake formation are mainly preserved within fine-grained lakebottom sediments or in buried delta deposits that are recorded in seismic profiles.
Delta deposition responded to minor and major accommodation changes across shorter and longer time scales. Short-term minor variations in accommodation space were probably related to lake-level changes in the range of a few metres, caused by seasonal or decadal changes in meltwater discharge and sediment supply (cf. Lønne & Nemec 2004;Gilbert & Crookshanks 2009). These changes mainly affected the accommodation space on the delta-brink zone, controlling the stability of the delta front and the related type of gravity-flow deposits on the delta slopes and delta bottomsets (Plink-Bj€ orklund & Steel 2004;Gobo et al. 2014;Talling 2014;Gobo et al. 2015;Dietrich et al. 2016;Hughes Clarke 2016). The larger-scale sediment dispersal pattern was controlled by the magnitude of major lake-level changes in the range of 20-60 m, the presence or absence of incised valleys and/or the number and depth of distributary channels and water depths (Dunne & Hempton 1984;Postma 1995;Muto & Steel 2001, 2004Uli cn y 2001;Ritchie et al. 2004a;Olariu & Bhattacharya 2006;Winsemann et al. 2009Winsemann et al. , 2011Eilertsen et al. 2011). The differences in sedimentary facies, thickness and slope angle of the foresets were controlled by the feeder system and accommodation available during these base-level changes (Postma 1995;; Uli cn y 2001; Sohn & Son 2004;Winsemann et al. 2009Winsemann et al. , 2011Eilertsen et al. 2011;Gobo et al. 2014Gobo et al. , 2015.
Deposition during lake-level rise. -During overall lakelevel rise an upslope shift of depocentres occurred. Vertically stacked shoal-water delta mouthbar deposits (FA2.1) formed on top of delta-plain deposits (FA1.1) during low rates of lake-level rises, when progressive aggradation of fluvial and/or delta-plain facies occurred in proximal areas (Posamentier & Morris 2000). These sandy mouthbar deposits display lake-ward dipping, laterally persistent low-angle foresets and form vertically stacked large-scale convex-up bedforms with good preservation of formsets (Fig. 9A, B), suggesting aggradation within increasing accommodation space in front of a retrograding shoal-water delta on a drowned glacifluvial delta plain Sohn & Son 2004;Fielding et al. 2005;Winsemann et al. 2009). In these transgressive systems the recurrence time of channel bifurcation and lobe switching of terminal distributary channels was long, allowing the channels to extend and accumulate as elongate sediment bodies (Figs 9A, B, 11;Olariu & Bhattacharya 2006). The seismic profile of the Porta fan and delta complex indicates that these shoal-water mouthbar deposits are genetically linked to high-angle Gilbert-type foresets that are exposed on the eastern margin of the truncated fan (Fig. 3A, seismic units 1-2). The sedimentary facies is dominated by deposits of supercritical surge-type turbidity flows (FA1.2.3). The small-scale fining-upward sequences of gravelly cyclic-step deposits and sandy antidune deposits (Fig. 6I-M) were probably triggered by frequent small-volume gravitational collapses of the upper delta slope (Talling 2014;Dietrich et al. 2016;Hughes Clarke 2016) during high rates of delta-front aggradation (Gobo et al. 2014(Gobo et al. , 2015. Metre-scale fining-and coarsening-upward trends may indicate seasonal or decennial variations in meltwater flows and a related fluctuation of the lake level and the delta-plain accommodation (Gobo et al. 2014(Gobo et al. , 2015. During high rates and magnitudes of lake-level rise, backstepping of delta lobes occurred, which decreased in thickness and lateral extent. The retrograding deposit profiles are stair-stepped (Fig. 3B, seismic units 1-4; Fig. 3C, seismic units 6-7), indicating a rapid upslope shift of depocentres (Muto & Steel 2001;Catuneanu et al. 2011;Villiers et al. 2013;Martini et al. 2017).
Deposition during highstand. -During lake-level highstand accommodation space progressively decreases and the stratal stacking pattern changes from aggradation to progradation with subhorizontal or falling delta-brink trajectories and an oblique erosional toplap geometry, which onlap the inherited depositional profile (Pore z bski & Steel 2006;Catuneanu et al. 2011). The sedimentation is characterized by thick high-angle foreset bedding, suggesting steep slopes of deep-water Gilbert-type deltas with gravity-driven flows (Nemec 1990;Postma 1995;Uli cn y 2001;Eilertsen et al. 2011;Gobo et al. 2015). The highstand Gilbert-type deltas comprise both coarsegrained gravelly or finer-grained sandy systems, depending on the feeder system.
The coarse-grained gravel-rich delta systems commonly show open-work gravel lenses in the lower delta slope and toeset area (FA1.2.2), indicating frequent slope-failure events and related cohesionless debrisflow and debris fall processes (Sohn et al. 1997;Nemec et al. 1999;Uli cn y 2001;Sohn & Son 2004). The alternation of steeply dipping coarse-grained foreset beds with abundant open-work gravel lenses and more gently dipping sandy foreset beds, deposited from more diluted flows (Fig. 6E-H) may point to autocyclic delta-slope steepening (Falk & Dorsey 1998;Longhitano 2008) or short variations in lake level and sediment supply, related to seasonal or decennial rates in meltwater production and sediment supply (Gilbert & Crookshanks 2009;Gobo et al. 2014Gobo et al. , 2015. The debris-fall dominated foreset facies assemblage would then record deposition during times of low-magnitude lake-level rise because the aggrading delta front then tends to store sediment and undergoes frequent gravitational collapses (Gobo et al. 2014(Gobo et al. , 2015. The absence of major bottomset deposits during this stage is related to the predominance of lowmobility debrisflows (Nemec 1990), whereas the turbidite-dominated facies assemblage would have formed Fig. 11. Characteristics of deltaic deposition, stratal stacking patterns and geomorphology under lake-level change. The geomorphological sketches compile data from this study, Muto & Steel (2001, 2004, Ritchie et al. (2004a, b), Olariu & Bhattacharya (2006), Lee et al. (2007), Winsemann et al. (2011) and Villiers et al. (2013). predominately during short periods of lake-level stillstand or slow fall, when the delta-front accommodation is at a minimum and sediment tends to be transported down-slope by erosional hyperpycnal flows (Gobo et al. 2014(Gobo et al. , 2015. The lateral transition into thick sandy bottomsets indicates the coeval deposition of turbidites in the delta-foot zone. Some intercalations of coarsergrained debrisflow deposits may point to deposition in front of delta-slope chutes, which transferred coarsergrained debrisflows to the delta-foot zone (Nemec 1990;Gobo et al. 2014).
The sand-rich highstand deltas are dominated by tractional bedforms, including climbing-ripple crosslamination, trough cross-stratification, low-angle crossstratification and backset cross-stratification, deposited by subcritical and supercritical turbidity currents (FA1.2.4-FA1.2.5). The grain size of the foreset-bed packages commonly decreased during progradation and a lateral facies transition from FA1.2.4 to FA1.2.5 can be observed. This might be related to an increase of the alluvial plain, increased sediment partitioning and the progressive deposition of finer-grained deposits on the delta slope (Posamentier & Morris 2000). Alternatively, the fining during progradation may be related to a decreasing water discharge and sediment supply and deposition from lower-energy density flows. Deposits of cyclic steps mainly occur within the coarsest foreset beds. In contrast to the cyclic-step deposits of the transgressive delta foresets (FA1.2.3) backsets are thicker, occur over the entire foreset length and show less variation in grain size, pointing to more sustained turbidity currents (Figs 7C-E, 8A, B). The finer-grained sandy foreset beds, deposited from migrating (humpback) dunes and ripples (Figs 7F-G, 8E) also require sustained turbidity currents that may reflect plunging hyperpycnal flows (Plink-Bj€ orklund & Steel 2004;Winsemann et al. 2007;Ghienne et al. 2010;Ventra et al. 2015;Carvalho & Vesely 2017) during low rates of delta-front aggradation. Gobo et al. (2015) suggested that a high proportion of foreset turbidites is related to a fairly persistent sediment bypass of the delta front when the delta brinkzone accommodation decreases or is at a minimum. The frequent occurrence of upslope migrating climbing ripples may indicate the zone of flow transition of plane-wall jets emerging from the delta-plain channels (cf. Jopling 1965;Clemmensen & Houmark-Nielsen 1981;Winsemann et al. 2007).
In contrast, Talling (2014) argued that thick tractional bedforms probably cannot be deposited by plunging hyperpycnal flows, as the suspended sediment concentrations in rivers are commonly too low, and they often do not coincide with flood peaks. However, Ghienne et al. (2010) were able to trace large climbingdune cross-stratified sandstones from the delta plain into upper foreset beds, clearly pointing to the existence of plunging sediment-laden, hyperpycnal meltwater flows.
The fine-grained foresets with climbing-ripple crosslaminated sand contain a few thick intercalations of pebbly sand beds deposited from cyclic steps (Fig. 8C,  D). These cyclic steps probably indicate infrequent larger slope failure events with longer run-outs, which may have been partly related to major flood peaks (Ventra et al. 2015). However, the exact trigger mechanism for the larger migrating supercritical bedforms remains uncertain because the foreset-topset transition zone is not preserved.
The chute fills were deposited from supercritical and subcritical turbidity currents, which may have resulted from the confinement of the currents (Gobo et al. 2015). High sedimentation rate under hydraulic-jump conditions are indicated by the formation of climbing humpback dunes with oversteepened and contorted foresets Lang & Winsemann 2013). Synsedimentary normal faults located at the channel margins seem to have favoured a vertical channel stacking (Winsemann et al. 2009).
The bottomsets (FA1.3.1) of the turbidite dominated sandy foresets are only poorly exposed. GPR profiles suggest that thick bottomset deposits are absent and foreset beds downlap prodelta deposits with an angular to tangential geometry, indicating rapid progradation.
Deposition during lake-level fall (forced regression) and lake-level lowstand. -During forced regression strong progradation took place. In seismic profiles forced regressive deposits are characterized by downstepping delta-brink trajectories. The tops of forced regressive deposits are either stepped-topped and attached or stepped-topped and detached (Figs 1, 3). Fluvial incision during downstepping led to erosion of the highstand deposits and the formation of deeply incised valleys and/or distributary channels. The large scale of cross-stratification and climbing-dune assemblages in the distributary-channel and incised-valley fills suggest high-gradient streams and high flow depths (cf. Massari & Parea 1990;Breda et al. 2007;Winsemann et al. 2011). Depositional processes on the delta slopes include debrisflows, subcritical to supercritical turbidity currents or tractional currents, depending on the remaining water depth, slope steepness and type of feeder system. Therefore, the type of forced regressive and lowstand foreset packages is highly variable and may range from highangle foreset bedding (FA1.2) to coarse-grained lowangle shoal-water mouthbar types (FA2.2) with reduced thickness (Fig. 11).
The development of stepped-topped detached Gilbert-type delta deposits was favoured by high magnitudes of lake-level fall, which promoted the development of incised valleys and the deposition of detached forced regressive coarse-grained delta lobes in front of the valleys , as shown in numerical simulations by Ritchie et al. (2004a).
Initial valley incision was probably caused by the formation of cyclic steps during rapid base-level fall (Strong & Paola 2008;Winsemann et al. 2011;Muto et al. 2012). Long-wavelength bedforms (FA1.1.2; Fig. 5) on the delta plain of the Emme delta, which formed during final lake drainage, provide evidence for incision by supercritical flows. Coeval sediments of the delta-foot zone may be represented by thick sand bedswith climbing humpback-dune stratification (FA1.3.1; Fig. 6C), which record high-energy turbulent waning flows under hydraulic-jump conditions at the mouth of the incised valley channel . The incised valleys captured the sediment and focussed the sediment supply to coarse-grained regressive lobes in front of the incised valley, leading to the development of digitate, tongue-shaped delta morphologies. These forced regressive deposits consist of sharp-based, high-angle foresetbed packages (30°-10°), up to 25 m thick. The height indicates that these forced regressive deltas prograded into relatively deep water. These foreset deposits are dominated by debrisflow deposits (FA1.2.1; Fig. 6A-C) that correspondwith strong fluvial erosion, a related high sediment supply to the delta front where debrisflows were deposited en masse when the slope diminished (Ilgar & Nemec 2005;Winsemann et al. 2011). The coarsegrained debrisflow dominated forced regressive delta lobes may be downlapped and partly overlain by smallerscale sandy delta lobes, deposited from debrisflows and turbidity currents (FA1.2.2) that record an upslope-shift of depocentres during valley back-filling. Valley filling mainly occurred during decreasing rates of lake-level fall and low base level (Blum & T€ ornqvist 2000;Winsemann et al. 2007Winsemann et al. , 2011Petter & Muto 2008). The formation of attached forced regressive delta deposits was favoured by a lower rate and magnitude of lake-level fall, a high rate of sediment supply and relatively steep slope gradients (Posamentier & Morris 2000;Ritchie et al. 2004a, b;Catuneanu et al. 2011), causing only minor incision in the upper portion of the deltas. The formation of relatively fixed, deep distributary channels, incised into the older delta plain and delta foresets, hindered major lateral delta-lobe shifting and led to the formation of various delta lobes that fringe and downlap the older delta body . The sedimentary facies offoresets and chute fills is dominated by tractional bedforms (FA1.2.3-FA1.2.4) with a high variability in grain size, probably deposited from plunging sedimentladen, hyperpycnal meltwater flows (cf. Ghienne et al. 2010;Gobo et al. 2014Gobo et al. , 2015. Coarsening-upward, prograding, shoal-water delta systems are indicative of forced regressive systems (Olariu & Bhattacharya 2006;Lee et al. 2007) where high-energy flows entered shallow water and a strong deceleration of the flow led to rapid deposition and aggradation of the sediment in the terminal distributary channel area (Dunne & Hempton 1984;Postma 1995;Uli cn y 2001;Sohn & Son 2004;Ilgar & Nemec 2005). Some channel bases show steep-flanked V-shaped scours and large intraclasts (Fig. 10D, E). These features are common in terminal distributary channels (Olariu & Bhattacharya 2006;Lee et al. 2007) and might be related to the formation of cyclic steps during base-level fall or major drainage events, when supercritical flow conditions established (Strong & Paola 2008;Muto et al. 2012). The presence of upstream accretion suggests that these lobes were quickly abandoned (Lee et al. 2007). The mound-shaped geometries and large range of palaeocurrent directions suggest a series of coalescing depositional lobes, expanding into the lake basin and creating a coarse-grained lower fringe, downlapping the older steeply dipping Gilbert-type delta foresets. Compared to transgressive mouthbar systems coarse-grained forced regressive shoal-water deltas have a larger number of terminal distributary channels and much shorter recurrence intervals of channel bifurcation, avulsion and lobe switching, resulting in an overall lobate shape ( Fig. 11; Olariu & Bhattacharya 2006;Lee et al. 2007). Coeval channel abandonment and decrease in the number of channels may occur in upper parts of the delta, leading to incision and increased discharge through the main distributary channels in the upper delta plain (Olariu & Bhattacharya 2006). In modern examples, commonly no major incision at the top of mouthbars has been observed and major incision is therefore regarded as indicative of base-level fall or major drainage events (Olariu & Bhattacharya 2006). However, the forced regressive coarse-grained mouthbar deposits of the study area are mainly vertically stacked, have partly well-preserved bar tops and are only slightly progradational (Figs 9C,D,10). This indicates that deposition took place into relatively 'deep' water (several metres) that provided sufficient accommodation space and prevented strong erosion and bypass.

Deformation structures
The deformation structures within the ice-marginal deltas comprise both contractional and extensional features, which are related to (i) gravitational tectonics; (ii) glacioteconics, (iii) crestal collapse above salt domes, and (iv) postglacial faulting during glacial-isostatic adjustment Winsemann et al. 2011). A neotectonic component cannot be ruled out in some cases (Brandes & Tanner 2012).
Gravitational tectonics. -Many river deltas show gravitational deformation that is expressed in a linked extensional and compressional fault system, where extension in the upper parts of the delta body (Figs 3B, 4B) is compensated by contraction at the delta-toe (Bilotti & Shaw 2005;Bini et al. 2007;Brandes et al. 2007aBrandes et al. , 2011. The extension in the upper part of the delta leads to the formation of basinward-dipping listric growth faults and half-graben structures. At the toe, compressional featu-res like thrusts and folds occur (King et al. 2010). The normal faults and the thrusts are commonly rooted in a basal detachment that links both structural regimes (Brandes et al. 2007b). Such a basal detachment is often controlled by overpressured shale (Cobbold et al. 2004) or major lithological contrasts ). Gravitational deformation is not restricted to large deltas at continental margins with a long life span, but also occur in small lake deltas (Bini et al. 2007;Brandes et al. 2011) and can be also reproduced by metre-scale analogue models (McClay et al. 1998(McClay et al. , 2003 and flume-tank experiments (Heller et al. 2001). In the study area, lithological contrasts between the Mesozoic bedrockand the overlying Pleistocene coarse-grained delta deposits may have supported the development of local detachments, which decoupled the gravitationally spreading delta body from the underlying bedrock. In some parts of the icemarginal deltas fine-grained lake-bottom sediments underlie the delta and may have supported the development of a basal detachment . The apparent absence of gravitational compression structures in the studied delta systems could be an effect of the limited upslope extension. Alternatively, these compressional structures are present but have not been recorded in the seismic profiles.
Glaciotectonic deformation. -Glaciotectonic deformation seems to play a minor role and only a few thrust sheets have been observed in sandy delta deposits. One reason might be that forced regressive delta systems commonly reflect lake drainage during an overall phase of ice retreat (Powell 1990;Ashley 1995;Lønne 1995;Winsemann et al. 2011;Girard et al. 2015). Additionally, the stable position of many deltas in front of bedrock highs prevented ice advance and related glaciotectonic deformation. However, the Elsterian Betheln delta must have been overridden by ice during the subsequent Saalian glaciation (Roskosch et al. 2015). A possible explanation could be an effective decoupling of the ice from the underlying sediments, which is often controlled by the water pressure at the ice/sediment interface (Kjaer et al. 2006), where an increase in water pressure can cause localized ice/bed decoupling (Fischer et al. 2011). An additional controlling factor for the deformation is the rheology of the material. The lack of deformation could be caused by the presence of frozen sediment, which is more stable (cf. Tylmann et al. 2012) and thus potentially less prone to deformation. Another option to explain the absence of glaciotectonic deformation structures is that some of the outcrops are probably too small to show these features, especially when the structures are large and the spacing between the individual thrust planes is high.
Deformation by dead-ice melting. -Normal faults in icemarginal deposits have often been regarded as diagnostic for the melting of dead ice in the subsurface (Selsing 1981;Prange 1995;Juschus 2001). Characteristic for dead-ice melting is a circular pattern with strongly curved faults in the sediments, reflecting the shrinking of the buried ice block and the related collapse of the hanging-wall material. Such a circular fault pattern can be observed around actively melting dead-ice blocks (Kjaer & Kr€ uger 2001) and can be also reproduced by analogue models that simulate depletion-related surface effects (Poppe et al. 2015). However, our field examples clearly indicate that dead-ice melting did not play a major role in the formation of normal faults.
Crestal collapse above salt domes. -The fill of the Central European Basin System is characterized by a large number of salt structures. Many of them reach close to the earth's surface and consequently, salt movements can have an impact on the Pleistocene sediments. Lang et al. (2014) showed in their modelling study that salt structures can be reactivated by ice-loading. An ice advance towards a salt structure causes salt flow from the source layer below the ice sheet towards the salt structure, resulting in uplift. When the diapir is overridden by the ice sheet the salt structure is pushed down. During ice retreat large parts of the displacement are compensated by a reversal of the salt flow, resulting in a renewed uplift. In such a setting, crestal collapse with normal faulting can be a common trigger for extensional deformation (Currie 1959;Alves et al. 2009). Comparable phenomena were shown by Lehn e & Sirocko (2005) in NW Germany and Al Hseinat et al. (2016) for the Baltic Sea, where faulting and surface subsidence is related to ongoing movements along local graben structures and the rise of salt diapirs.
Postglacial faulting during glacial-isostatic adjustment. -Glacial-isostatic adjustment can lead to the reactivation of pre-existing faults in the subsurface due to lithospheric stress field changes as a consequence of the growth and decay of large ice sheets (Kukkonen et al. 2010). Pre-existing faults in the basement can be reactivated and propagate into the overlying sediments that sealed the tip lines of the faults . Seismic profiles of the Emme delta (Fig. 3B) and the Porta fan and delta complex (Fig. 3A) show normal fault systems developed in Mesozoic rocks, which can be traced into the overlying Pleistocene sediments. The reactivation of the Mesozoic normal faults in this location is interpreted as a consequence of extension in the forebulge area of the advancing ice sheet, in combination with loading by a glacial lake . The fault activity ceased after the lake had considerably drained, probably indicating that fault activity in these cases was controlled by water load and water pressure and that below a critical threshold fault activity ceased . The dense arrays of shear-deformation bands (Fig. 8A, B), which are developed within the Freden delta, also formed above the tip line of buried Mesozoic faults and therefore most likely indicate a fault reactivation due to lithospheric stress changes caused by glacial-isostatic adjustment during MIS 8 (Brandes et al. 2018).

Conclusions
The studied forced regressive ice-marginal deltas are considered as representative of delta styles in glacial lake basins, affected by rapid base-level fall. They have many characteristics in common with other (glacigenic) Gilbert-type and shoal-water mouthbar type deltas, including the stair-stepped fan, lobate or more digitate tongue-shape geomorphology, the large-scale depositional architecture and range of sedimentary facies.
The frequent occurrence of bedforms deposited by supercritical turbidity flows along the foresets may be a characteristic feature of high-energy ice-marginal deltas. Bedforms typically comprise laterally and vertically stacked successions of cyclic steps and antidunes. Trigger mechanisms of supercritical flows were hyperpycnal meltwater flows and slope-failure events in response to accommodation changes on the delta plain.
Delta deposition responded to minor and major accommodation changes across shorter and longer time scales. Short-term minor variations in accommodation space were probably related to lake-level changes in the range of a few metres, caused by seasonal or decadal changes in meltwater discharge and sediment supply. These changes mainly affected the accommodation space on the delta-brink zone, controlling the stability of the delta front and the related type of gravity flow deposits. Surge-type (supercritical) turbidity currents were probably triggered by small-volume gravitational collapses of the upper delta slope during periods of slow lake-level rise when high rates of delta-front aggradation occurred. In contrast, more sustained (supercritical) turbidity currents were probably triggered during lakelevel highstand and lowstand by hyperpycnal plunging meltwater flows, when accommodation space on the delta plain was low.
The larger-scale depositional delta architecture was controlled by the magnitude and rate of major lake-level changes. The differences in sedimentary facies, thickness and slope angle of the foresets during forced regression were controlled by water depth, the presence or absence of incised valleys and/or the number and depth of distributary channels. Incised valleys formed during high-magnitude lake-level falls. These deep valleys focussed the sediment supply to coarse-grained elongate, tongue-shaped lobes, which were mainly deposited by cohesionless debrisflows. During lower magnitudes of lake-level fall or high-magnitude falls with high sediment supply attached sand-rich forced regressive aprons formed. If water depths became very low coarse-grained shoal-water mouthbar deltas formed that fringe and downlap the older Gilbert-type deltas.
The exposed delta sediments mainly comprise highstand, forced regressive and lowstand deposits, which record the phase of maximum lake level and subsequent successive lake drainage. The stair-stepped profiles of the delta systems reflect the progressive basinward lobe deposition during forced regression when the lakes successively drained. Deposits of the stair-stepped transgressive delta systems are buried and downlapped by the younger forced-regressive deposits and only deposits of the maximum lake-level highstand are preserved in geomorphology, forming the uppermost unit of the delta systems. The good preservation of delta sediments during overall lake-level fall is related to the deglaciation stage, during which meltwater volumes and sediment supply are high. Depending on the rate and magnitude of lake-level fall trumpet-shaped deeply incised valleys or a larger number of deep distributary channels formed, leading to telescoping, tongue-shaped, fan-shaped or lobate Gilberttype delta morphologies. Foreced-regressive shoal-water mouthbar deltas are typically lobate, which resulted from the coalescence of multiple terminal distributary channels and mouthbars. In contrast, shoal-water deltas deposited during transgression have more stable channels and the recurrence time for channel bifurcation and lobe switching is long, allowing the channels to extend and accumulate as elongate sediment bodies with a more digitate tongueshape.
Deformation structures within the ice-marginal deltas comprise both contractional and extensional features, which are related to (i) gravitational delta tectonics; (ii) glaciotectonics, (iii) crestal collapse above salt domes and (iv) postglacial faulting during glacial-isostatic adjustment. In some cases, a neotectonic component cannot be ruled out. Dead-ice melting did not play a major role in the formation of normal faults.