Zackenberg is located on the Wollaston Forland in northeast Greenland (74∘30′ N, 20∘30′ E; Fig. 1), 90 km east of the Greenland Ice Sheet margin. This region was glaciated several times during the Quaternary period (Hjort, 1981; Funder et al., 1994; Bennike et al., 2008). The last glaciation culminated in the LGM during the late Weichselian, when the Greenland Ice Sheet extended across the region (Bennike et al., 2008). Sedimentary archives and the geomorphology of the continental shelf and slope indicate that during the LGM, ca. 22 ka, grounded glacier ice extended onto the outer continental shelf (Evans et al., 2002; Ó Cofaigh et al., 2004; Winkelmann et al., 2010). Streamlined subglacial bedforms record the presence of erosive, warm-based ice streams in the major fjords and cross-shelf troughs in NE Greenland during this period. Intervening areas were covered by grounded but non-streaming ice, in some cases preserving pre-existing landforms (Funder et al., 2011). Deglaciation initiated in the coastal areas east of Young Sound between 10.1 and 11.7 ka (Bennike et al., 2008). The inner-fjord areas were ice-free by between 9.5 and 7.5 ka. At Zackenberg, the earliest postglacial radiocarbon date from a marine fossil is 10.1 ka, suggesting the study area was ice-free by this time (Christiansen et al., 2002).

Following deglaciation, the sea inundated low-lying areas of the Zackenberg Valley. Rapid emergence is documented during the early Holocene due to postglacial crustal rebound. Regional RSL curves are reconstructed from the height of raised beach deposits using optically stimulated luminescence (OSL) dating and AMS 14C dating of palaeosurfaces and marine fossils (Bennike and Weidick, 2001; Pedersen et al., 2011). Christiansen et al. (2002) reconstructed RSL changes at Zackenberg using a combination of geomorphological evidence and radiocarbon dates of marine macrofossils and terrestrial plant remains. This curve indicates a late Quaternary marine limit of ca. 40 m a.s.l. During the early Holocene, RSL declined rapidly, reaching present sea level by ca. 5–6 ka (Christiansen et al., 2002).

The erosional competence of ice streams during the LGM, deglaciation history, and subsequent changes in RSL have important implications for both landscape and permafrost development in Zackenberg. Erosive ice streams in the fjords removed much of the sedimentary record from previous glacial–interglacial cycles. At the same time, warm-based glaciers precluded the formation or preservation of permafrost due to frictional heat generated at the base of the sliding glacier and the trapping of geothermal heat under the ice (Humlum, 2005). The age of permafrost in the Zackenberg lowlands is likely linked to the timing of regression, as warm boundary conditions at the sea floor would have prevented permafrost aggradation.

Mean annual air temperature at Zackenberg between 1996 and 2013 was -9.0 ∘C. The mean annual precipitation for the same period was 219 mm water equivalent, of which 90 % fell as snow or sleet. Sea ice covers the Young Sound between 9 and 10 months each year from late October (Hansen et al., 2008; Jensen et al., 2014). Hydrology and sediment transport in the Zackenberg River have been summarized by Hasholt et al. (2008). The drainage basin covers an area of 512 km2, of which 20 % is glaciated. Runoff in the river typically begins in June and continues until September. Sediment transport is dominated by extreme discharge events. In addition to the Zackenberg River, a number of minor rivers and streams drain the hillslopes and lowland area.

Zackenberg is located within the continuous permafrost zone. The permafrost thickness is modelled to be between 200 and 400 m (Christiansen et al., 2008). Ground temperatures are monitored to a depth of 20 m at two locations within the study area (C1 and C2 in Fig. 1c). Temperatures at 20 m depth are ca. -6 ∘C, with little interannual variation since monitoring began in 2012. Seasonal thaw progression and active-layer thickness has been measured since 1996 (Christiansen et al., 2003). On average, the active-layer thickness varies between 70 and 80 cm – with topography, and its impact on the snow regime and hydrology, being the primary controlling factor (Christiansen et al., 2008).

The Zackenberg Valley is oriented along a NW-trending fault (Koch and Haller, 1971; Nøhr-Hansen et al., 2011), separating Caledonian gneiss and granite exposed to the west in the Zackenberg Mountain (1303 m a.s.l.) from Cretaceous and Jurassic sedimentary rocks exposed to the east in the Aucellabjerg Mountain (911 m a.s.l.). The valley is ca. 8 km long, expanding from 2 to 7 km in width towards Young Sound (Fig. 1c). The lowlands are defined as the portion of the landscape below the early Holocene marine limit (40 m a.s.l.). The lowlands are primarily a relict landscape, formed by glacial, fluvial, marine, and periglacial processes during glaciation, deglaciation, and RSL decline. Glacial and glaciofluvial landforms include moraine ridges, ground moraine, meltwater plains, and raised delta terraces (Christiansen and Humlum, 1993; Christiansen et al., 2002). Permafrost landforms include ice-wedge polygons and palsas (Christiansen, 1998a; Christiansen et al., 2008).

The raised palaeo-delta at the mouth of the Zackenberg Valley covers an area of ca. 2 km2 and consists of a series of terraces. To the west, the spatial extent of the palaeo-delta is restricted by the Zackenberg Mountain. The eastern boundary is less easily defined and grades into glacial and glaciofluvial deposits. Christiansen et al. (2002) determined that the deltaic deposits began to form prior to 9.5 ka.

FA I was observed at all sites except those bordering the present-day shoreline (S1–S5). Exposures of FA I ranged up to 6 m in thickness and consisted of a compact, sandy, matrix-supported diamicton (facies 1). The diamicton contained angular to subrounded clasts, up to boulder size, of varying provenance. The deposits were overconsolidated when compared to other sedimentary facies. The lower boundary was not observed. The upper boundary to FA II was sharp and decreases in elevation, from ca. 41 m in the upper part of the delta to ca. 12 m towards the modern-day coastline (Fig. 5).

FA I is interpreted as a subglacial till on the basis of its stratigraphic position, compaction, lithology, and lateral extent (Evans and Benn, 2004). The decline in elevation of FA I towards the southeast likely reflects the increasing depth to bedrock. Widespread till deposits have previously been identified in the Zackenberg Valley bottom by Christiansen and Humlum (1993) and Christiansen et al. (2002) and outcrop at the ground surface in the glacial and glaciofluvial lowlands to the east of the delta terraces (Fig. 1). FA I was likely deposited during the last glaciation, when the Zackenberg lowlands were covered by an ice stream.

FA II includes silts (facies 2), interbedded sands and silts (facies 3), and normally graded sands and silts (facies 4). Laminated sands (facies 6) and massive sands (facies 7) are recorded in the lower portion of FA II at S6 and S7. FA II outcrops at all locations with the exception of S3, S4, and S9 (Fig. 5). Individual beds appear horizontal. The thickness of this unit ranged between 2 and 10 m. The upper boundary is transitional when overlain by FA III and sharp and erosional when overlain by FA IV. At S6, C1, and C2, FA II initially fines upwards before an upwards coarsening into FA III (Fig. 5). At these sites, coarse-grained facies (facies 6 and 7) were observed at the base of FA II. In other sections, FA II generally fines upwards. Bioturbation was uncommon and restricted to facies 2. Macrofossils consisted of bivalve shells and shell fragments. Isolated outsized clasts were observed.

FA II is polygenetic and was deposited into a fjord environment during RSL highstand during and following deglaciation. The deposits consist of parallel-bedded fine sands and silts and reflect deposition by suspension fallout from hyperpycnal plumes and low-density turbidity currents (Table 1). The upwards fining and thinning observed at S6 and in C1 and C2 records the withdrawal of the sediment source as the ice front retreated north during deglaciation towards its grounding line at the mouth of the Zackenberg Valley (Ó Cofaigh et al., 1999). The upper portion of FA II is characterized by an upwards coarsening and thickening. This unit is interpreted as the prodelta environment. The high sand content of deposits from suspension is typical of steep, shallow-water deltas, where sand is carried beyond the delta slope by hypopycnal plumes (Corner, 2006). Turbidites record instabilities on the surrounding slopes, periods of high-river discharge, variations in glacier front position, or collapsing sediment plumes (Gilbert, 1983; Hansen, 2004). The relative absence of bioturbation and limited number of trace-making organisms indicates a stressed environment with turbid water conditions and relatively high sedimentation rates (Netto et al., 2012). Overall, the vertical changes in the relative dominance of the facies records an approaching delta front.

FA III consists of interbedded sands and silts (facies 3), graded sands and silts (facies 4), laminated sands (facies 6), massive pebbly sands (facies 7), and stratified pebbly sands and gravels (facies 8). FA III is exposed at S2–S6 and in both C1 and C2. Exposures range in thickness from 2 to 10 m (Fig. 5). The relative influence of facies 3 and facies 4 decreases upwards as the unit coarsens and beds increase in thickness. Evidence for soft-sediment deformation is occasionally observed in association with dewatering structures. At sites S2–S6, beds dip 5–25∘ southeast, towards Young Sound (Fig. 4a). Dip angles increase upwards as the beds coarsen and thicken towards the top of the exposures.

The inclined, planar beds of FA III are interpreted as foresets deposited during the progradation of a Gilbert-type delta. The overall decline in the elevation towards Young Sound indicates that this facies association was deposited during RSL fall. Variations in gain size and facies composition reflect differences in their proximity to fluvial channels, variations in discharge, and sediment transport processes. The deposits reflect suspension settling, turbidity currents, debris flows, and grain flows associated with slope failure on the delta front. Facies 7 and 8 were deposited by gravitational avalanches and debris flows following accumulation on the upper delta slope (Nemec, 1990; Plink-Björklund and Ronnert, 1999; Hansen, 2004). Facies 3, 4, and 6 record the activity of turbidity currents either due to underflow of sediment-laden river water or flow transformation (Kneller and Buckee, 2000; Winsemann et al., 2007). The presence of soft-sediment deformation structures and absence of bioturbation suggest a rapid sedimentation rate for FA III.

FA IV consists of cross-stratified sands (facies 5), laminated sands (facies 6), massive pebbly sands (facies 7), and gravels (facies 9). The facies association ranges between 0.5 and 3.0 m in thickness and is exposed at elevations between 5 and 45 m a.s.l. The elevation of FA IV declines towards Young Sound both to the south and east (Fig. 5). Facies 9 increases in dominance with distance inland. FA IV has a sharp, erosive contact with the underlying unit.

FA IV is polygenetic and the component facies are related to a range of depositional processes and environments. Deposits are interpreted to primarily reflect fluvial activity associated with a deltaic distributary plain or braided-river system. Diffusely stratified gravels with imbrication (facies 9) and planar-parallel laminated sands (facies 6) indicate bedload transport by unidirectional currents of varying strength (Miall, 2010). Cross-stratified sand beds record the migration of dunes in braided-river channels or distributary systems (Miall, 2006; Winsemann et al., 2007). Massive pebbly sands (facies 7) are deposits of pseudoplastic debris flows (Miall, 2010). Secondary processes include winnowing by wind, marine reworking during RSL fall, and redistribution of sediment by snowmelt (Christiansen and Humlum, 1993; Christiansen, 1998b; Christiansen et al., 2002). FA IV developed at successively lower levels during RSL fall and fluvial incision and channel migration.

Few studies have examined the relationship between ground-ice characteristics and postglacial landscape development in Greenland (Gilbert et al., 2016). Pollard and Bell (1998), however, presented a model for ground-ice aggradation within epigenetic permafrost below the Holocene marine limit in the Eureka Sound lowlands on the Fosheim Peninsula in high-Arctic Canada. Here, the nature and distribution of ground ice was related to RSL change, sediment grain size, and epigenetic permafrost aggradation. On the Fosheim Peninsula, massive intrasedimental ice bodies are observed at the contact between marine silts and underlying gravels and sands, deposited during transgression. Low hydraulic conductivities in the overlying fine-grained sediments did not permit sufficient water migration resulting in the formation of reticulate, layered, lenticular, and pore cryofacies (Pollard, 2000a). Massive ice bodies and cryofacies were interpreted to have formed during and following deglaciation, when glacier meltwater and brackish sea water were available to recharge local groundwater systems (Pollard, 2000b).

C1 and C2 are generally ice poor and the majority of the cores are characterized by the pore cryofacies. Cryofacies with visible ice (La and Su) are restricted to the fjord-basin and delta-slope facies associations. A similar pattern is observed in gravimetric moisture content and excess ice content (Figs. 7, 8). The distribution of ground ice in the Zackenberg River delta is similar to that described by Pollard (2000a, b). The presence of the layered cryofacies in epigenetic permafrost has also been described by Kanevskiy et al. (2014) in lacustrine sediments in interior Alaska. The suspended cryofacies is likely a more developed form of the layered cryofacies, forming during increased moisture availability or slower rates of freezing. The absence of appreciable ground ice in the remainder of the core indicates that permafrost in the Zackenberg lowlands is primarily epigenetic and formed following sea-level fall during the early Holocene. The differences in ice content and cryofacies between C1 and C2 likely relate to either variations in site-specific moisture availability during permafrost aggradation or lateral variations in sediment characteristics that are unresolvable at the core scale.

The presence of syngenetic permafrost is inferred using a combination of sedimentology and the dating results in the top of C2 – which record continued sedimentation following subaerial exposure. Here, the development of syngenetic permafrost is related to localized accumulation of niveo-aeolian sediments and organic material (Christiansen, 1998b; Christiansen et al., 2002). The uppermost syngenetic component in FA IV is almost 3 m in thickness. Despite continuous sedimentation and sufficient moisture availability, syngenetic permafrost at C2 consists of pore cryofacies. This is attributed to the coarse-grained sediment and the absence of frost-susceptible silts. Syngenetic permafrost is thus of local occurrence on the palaeo-delta surface and relates to periglacial, nival, fluvial, and aeolian activity during the Holocene.

During the past 13 ka, the Zackenberg lowlands have undergone major environmental changes through periods of glacial, marine, fluvial, and periglacial dominance. Sedimentological, cryostratigraphic, and geomorphological observations are combined with the OSL dating results to produce a model for the landscape and permafrost development of the Zackenberg lowlands. This model is condensed into three stages – illustrating the major changes in sediment supply, sea level, and sedimentary processes following glaciation (Fig. 9). The model provides a framework in which to discuss changes in the sedimentary facies and cryostratigraphy and improves our understanding of the amount of ground ice in Arctic valleys. This is important as the amount of ground ice determines the potential landscape response due to climate change.

A stratigraphic model for fjord-valley infilling in Norway has been presented by Corner (2006). This tripartite model consists of a deglacial transgressive systems tract (DTST), a deglacial highstand systems tract (DHST), and a postglacial forced-regressive systems tract (PRST). The DTST is deposited during deglaciation by retrogradational stacking of proglacial sediments. Sedimentation is concurrent with marine inundation and the early stages of sea-level fall during isostatic rebound. The DHST is characterized by the rapid progradation of glaciofluvial deltas and concurrent deposition of glaciomarine deposits on the basin floor. Lastly, the PRST forms following areal deglaciation and is characterized by fluvial delta progradation, fluvial incision, and terracing during RSL fall. A similar model was developed for the Scoresby Sound region, ca. 750 km south of the Zackenberg area, by Hansen (2004). Though the details of the stratigraphy vary between locations, the underlying controls identified in these models – sediment supply, accommodation space, and RSL change – remain the same.

Recent studies have presented models for infilling of formerly glaciated fjord valleys from Greenland (Hansen, 2001, 2004), Norway (Nemec et al., 1999; Corner, 2006; Eilertsen et al., 2006; Hansen et al., 2009; Eilertsen et al., 2011), Canada (Marchand et al., 2013), Sweden (Plink-Björklund and Ronnert, 1999), and Germany (Winsemann et al., 2007). These investigations demonstrate that fjord valleys are primarily infilled during deglaciation and RSL fall, resulting in a complex stratigraphy, which can vary substantially between locations. However, recent studies also indicate that the underlying depositional regime controls are the same (Eilertsen et al., 2011). With the exception of the investigations in Greenland, these studies are restricted to non-permafrost environments. No previous studies have investigated ground ice or cryofacies in valley-fill deposits.

Previous studies have identified similar infilling patterns of valleys following regional deglaciation. Given similarities in sedimentary facies and landscape development, it is reasonable to expect that the aggradational history of permafrost is similar as well. This suggests that permafrost below the upper marine limit of large tributaries valleys throughout northeast Greenland is likely a Holocene phenomenon, with the exact age corresponding to the timing of subaerial exposure following marine transgression. These coarse-grained deposits contain an assemblage of cryofacies characteristics of ice-poor epigenetic permafrost.