Harald Moltke Bræ is an outlet glacier in northwest Greenland that is 6.3 km wide at its grounded terminus (Koch, 1928; Wright, 1939; Davies and Krinsley, 1962; Rignot et al., 2001; Figure 2). Early observations suggested that the terminus had two calving lobes (Wright, 1939) shown in Figure 3B. However, recent imagery shows this is no longer the case (Figure 3B). The catchment area of Harald Moltke Bræ is smaller than most other glaciers in the study area, draining only 1400 km² (Table 1).

Early studies provided a detailed historical account of terminus change for the period 1916–1965 (Wright, 1939; Davies and Krinsley, 1962; Mock, 1966; Figure 3B). Between 1916 and 1926 the glacier retreated, which was followed by advance until 1932 when it reached a similar position as in 1916 (Wright, 1939; Davies and Krinsley, 1962; Mock, 1966). Between 1932 and 1937, the calving of large tabular icebergs was observed (Wright, 1939) and, by 1959, the terminus had retreated 5.5 km from its 1932 position (Davies and Krinsley, 1962). It is also known that the glacier underwent net retreat of 2.9 km between 2000 and 2010 (Murray et al., 2015).

From 1916 to 1965 the velocity of Harald Moltke Bræ fluctuated greatly between 30 and 1000 m a⁻¹ (Mock, 1966). More recently, in 2000/01, the glacier was flowing at 30–100 m a⁻¹ at the terminus, reaching up to a maximum of 300 m a⁻¹ further up glacier (Joughin et al., 2010a). By 2005, the velocity at the terminus increased to 2000 m a⁻¹ (Joughin et al., 2010a). Several authors have suggested these recent changes in velocity reflect surge behavior at Harald Moltke Bræ (Rignot and Kanagaratnam, 2006; Moon et al., 2012). In particular, velocity increase during 2005 coincided with terminus advance (1.2 km) in 2004/2005 (Murray et al., 2015).

Since the detailed work of Mock (1966), few studies have specifically focused on Harald Moltke Bræ. Its surge-like behavior is unusual in comparison to nearby outlet glaciers. Despite showing marked increases in velocity and advance during proposed surge events (Rignot and Kanagaratnam, 2006; Joughin et al., 2010a; Murray et al., 2015), the glacier has undergone large net retreat following surge events (Murray et al., 2015). This has led to a retreat of ~12 km compared to its position in 1916 (Figure 3B) and hints that the glacier has been influenced by longer-term external environmental drivers.

Heilprin and Tracy Glaciers are two large outlet glaciers in northwest Greenland that are 6.7 and 5 km wide, respectively (Table 1), and collectively drain an area of ~12,000 km² into Inglefield Bay (Figure 3A). Davies and Krinsley (1962) noted that both glaciers had floating ice tongues in 1892, although early observations with limited data make this difficult to verify. Rignot et al. (2001) later found floating sections absent from both glaciers, which recent grounding line data confirms (Morlighem et al., 2014; Figure 2A). Basal topographic data from these glaciers (Figure 2B), shows they both lie below sea level for a distance of 36–42 km inland of the terminus (Morlighem et al., 2014).

Both Heilprin and Tracy Glaciers have undergone net retreat in the twentieth Century (Kollmeyer, 1980). Tracy Glacier retreated by 7 km between 1892 and 1959 (Davies and Krinsley, 1962) and between 1949 and 2009 the glacier retreated a further 15 km away from Josephine Peary Øer Island (Figure 3C; Dawes and van As, 2010). In contrast, Heilprin Glacier, retreated only 4 km between 1892 and 2009 (Porter et al., 2014). The retreat of Tracy Glacier has increased over the last two decades: between 2000 and 2005, Tracy Glacier lost 8 km of its terminus (Figures 4A,B) followed by 2 km retreat between 2005 and 2013 (Porter et al., 2014). These differing rates of retreat were attributed at least partly to fjord geometry, and primarily deeper basal topography below Tracy Glacier which could allow warm water intrusion (Porter et al., 2014). Surprisingly, despite an inland-sloping bed at Heilprin, it is undergoing slower dynamic change than Tracy Glacier (Porter et al., 2014).

Estimates of ice discharge from European Remote Sensing (ERS) data in the mid-1990s at these glaciers are given in Table 1. Porter et al. (2014) recorded a doubling in thinning rates between 2011 and 2012 at Tracy Glacier. In terms of ice velocity fluctuations, both Heilprin and Tracy Glaciers experienced increases in speed between 2000/01 and 2005/06 of 20 and 40%, respectively (Joughin et al., 2010a), and coincided with the large retreat observed at Tracy Glacier (Figure 4). There is no record of surging at these glaciers.

Despite retreat taking place at both glaciers, and aside from study by Porter et al. (2014), Tracy and Heilprin Glaciers have been subject to little in-depth research in comparison to other northern Greenland areas. The differing responses of these neighboring glaciers, potentially attributed to fjord geometry, suggests uncertainty remains regarding the controls on these outlet glaciers.

The north NW region (Figure 1) consists of two of the largest outlet glaciers in Greenland: Petermann and Humboldt. These glaciers have received the most recent research attention in northern Greenland.

The northernmost region of Greenland in our study (North Central, Figure 1) consists of eight marine-terminating outlet glaciers. The region was first documented in studies by early explorers (Peary, 1892; Rasmussen, 1919), but the majority of these glaciers have had little scientific attention in recent years.

At the far north-eastern margin of the GrIS lie several major outlet glaciers (Figure 1) which are Academy, Marie Sophie and Hagen Bræ Glaciers.

Nioghalvfjerdsfjorden and Zachariae Isstrøm are the two main outlets of the NEGIS (Khan et al., 2014) and collectively drain around 30% of the northern GrIS by area (Rignot and Kanagaratnam, 2006; Table 1). Nioghalvfjerdsfjorden is 24 km wide, and has an extensive (~69 km long) floating ice tongue (Figure 2A), that widens down-fjord to 30 km at the terminus (Thomsen et al., 1997). Zachariae Isstrøm is 27 km wide and terminates in an embayment typically surrounded by calved icebergs (Box and Decker, 2011). Zachariae Isstrøm previously terminated as a floating ice tongue (Figure 2A), but this dramatically disintegrated between 2000 and 2006, meaning that the glacier terminus is currently grounded (Khan et al., 2014). Both glaciers lie above deep basal troughs, which rest significantly below sea level (Mayer et al., 2000), and have a reverse bed slope (Bamber et al., 2013; Figure 2B).

Information on the frontal positions of these two glaciers has been comparatively limited due to year-round ice mélange in the fjords, resulting in an ambiguous calving region and therefore making it difficult to accurately identify the true calving front (Bevan et al., 2012; Murray et al., 2015). In 1976, the two glaciers were thought to have coalescing ice tongues (Weidick et al., 1995), which suggests they have since retreated substantially and become separated (Rignot et al., 2001). Early records from Nioghalvfjerdsfjorden found 20 km of retreat between 1950 and 1963 (Thomsen et al., 1997). Subsequently, the terminus retreated 5–7 km during 1978–2003 (Khan et al., 2014). Zachariae Isstrøm, lost 1400 km² of its ice shelf between 2000 and 2006 (Moon and Joughin, 2008) and its grounding line has begun to rapidly retreat downslope (Mouginot et al., 2015). Both glaciers underwent similar ice loss in 2004/05, with 60 km² of ice lost at Nioghalvfjerdsfjorden and 67 km² at Zachariae Isstrøm (Box and Decker, 2011). Two years later, during 2006/07, both glaciers advanced, although of differing magnitudes: 8.4 km² at Nioghalvfjerdsfjorden and 45.8 km² at Zachariae Isstrøm (Box and Decker, 2011). Subsequently, between 2009 and 2012, sections of the Nioghalvfjerdsfjorden retreated by a further 2–3 km (Khan et al., 2014).

Nioghalvfjerdsfjorden and Zachariae Isstrøm were considered slow-moving in the early 1950s (Helk and Dunbar, 1953), although no specific velocity values were given. However, more recent data show that they are relatively fast-flowing reaching speeds of >1 km a⁻¹ at their termini (Khan et al., 2014; Figure 6). No significant acceleration or deceleration was detected between 2000 and 2006 on Nioghalvfjerdsfjorden (Rignot and Kanagaratnam, 2006; Joughin et al., 2010a). However, a more recent study observed an acceleration of 100 m a⁻¹ between 2000 and 2011 (Khan et al., 2014). Zachariae Isstrøm has been accelerating since the early 2000s, increasing by up to 200 m a⁻², following the disintegration of part of the ice shelf in 2004/05 (Rignot and Kanagaratnam, 2006; Moon and Joughin, 2008; Joughin et al., 2010a). Acceleration then continued, although more steadily, until 2012 (Khan et al., 2014; Mouginot et al., 2015). Since 2012, ice velocities on Zachariae Isstrøm have increased by 25%, accompanied by accelerated frontal retreat (Mouginot et al., 2015).

Estimates suggest that these glaciers must have experienced submarine melt rates of between 6 and 8 m a⁻¹ to explain their 1996 ice flux (Rignot et al., 1997). This is similar to submarine melt rates recorded at other northern Greenland outlet glaciers e.g., Petermann Glacier (Rignot and Steffen, 2008). The break-up of fast ice offshore of Nioghalvfjerdsfjorden in 1997 was followed by a large glacier calving event and is thus considered an important control on the rates of calving (Reeh et al., 2001).

More generally, the NEGIS is thought to be undergoing dynamic thinning as a result of climate change, losing mass at a rate of >10 Gt yr⁻¹ (Khan et al., 2014). As the NEGIS extends far into the ice sheet interior (Khan et al., 2014; Figure 2C) and sits on a reverse bed slope, these glaciers have the potential to discharge large volumes of ice. Should retreat continue at the present or increased rates (particularly at Zachariae Isstrøm; Khan et al., 2014), subsequent ice flow speed-ups could cause significant mass loss from a large inland area of the GrIS (Csatho et al., 2014). This could further destabilize this region and increase its contribution to twenty first century sea level rise.

Storstrømmen is another large outlet of the NEGIS that, along with Kofoed-Hansen Bræ, drains ~120,000 km² (Table 1). The catchment extends to the summit of the GrIS (Reeh et al., 2003; Figure 2C). Kofoed-Hansen Bræ is the northeastern branch of Storstrømmen which discharges ~25% of the Storstrømmen ice flux (Mohr et al., 1998). Storstrømmen has a two-lobed calving front, one of which drains directly into the ocean and the other of which joins with L. Bistrup Bræ to the south (Figure 6; Khan et al., 2014). L. Bistrup Bræ terminates alongside Storstrømmen and is ~11 km wide with a catchment of ~20,000 km² (Table 1). Within this embayment, both Storstrømmen and L. Bistrup Bræ have floating ice sections (Figure 2A) that are 8.4 and 6.2 km long, respectively, whereas data suggests Kofoed-Hansen Bræ is grounded (Morlighem et al., 2015).

Relatively little is known about frontal position changes at these glaciers. Available data show that all three glaciers retreated between 2001 and 2005 (Seale et al. (2011). In total, this resulted in retreat of 0.3 km at Storstrømmen, and 0.76 km at both Kofoed-Hansen Bræ and L. Bistrup Bræ (Seale et al., 2011). However, during this period Storstrømmen underwent a short term advance of 0.4 km² in 2001/02 (Box and Decker, 2011). Later, between 2005 and 2008, only Storstrømmen glacier continued to retreat, while L. Bistrup Bræ and Kofoed-Hansen Bræ both advanced by 0.29 km (Seale et al., 2011).

The velocity of Storstrømmen was thought to be ~230 m a⁻¹ in 1996 (Rignot et al., 2001). L. Bistrup Bræ is considered slow moving, although specific values were not documented in the literature (Rignot et al., 2001), with almost zero velocity between 2000/01 and 2009/10 at the terminus (Joughin et al., 2010a). However, further inland the glacier flows at ~100 m a⁻¹ in 2008/09 (Figure 6). During this same period (2000/01–2009/10) Storstrømmen Glacier decelerated by 60 m a⁻² (Joughin et al., 2010a).

Evidence of surging exists for all three glaciers. A surge was documented at Storstrømmen between 1978 and 1984 (Reeh et al., 1994), when the glacier advanced ~12 km. By the 1990s it was considered to be quiescent (Mohr et al., 1998), and has remained so since that time (Reeh et al., 2003; Rignot and Kanagaratnam, 2006). L. Bistrup Bræ is also likely to be a surge-type glacier, based on thickening at the grounding line, although there is no direct evidence of an actual surge event. The glacier is instead hypothesized to have surged in the past and now be quiescent (Rignot et al., 2001). Both L. Bistrup Bræ and Storstrømmen have experienced very low flow speeds over the past decade, potentially indicative of quiescence (Moon et al., 2012), as well as a characteristic pattern of thickening inland, where the surge may have initiated, and thinning toward the terminus (Abdalati et al., 2001; Csatho et al., 2014). Surveys between 1995 and 1999 found thinning rates of 2 m a⁻¹ in the lower reaches of both glaciers and thickening up to 3 m a⁻¹ further inland (Thomas et al. (2009). Following this, however, thickening rates on the upper reaches of Storstrømmen began to decrease (Thomas et al., 2009).

Despite these three outlet glaciers collectively draining an area (27%) of the GrIS, which is far greater than the well-studied Petermann Glacier (Table 1), little previous work has focused on their dynamics. Considerable variability in terminus positions have taken place over the last two decades, particularly at Storstrømmen and L. Bistrup Bræ, which has been linked to surge activity. These glaciers are important in terms of draining a large proportion of the NEGIS and it is necessary for further work to better understand their surge-nature and their implications for increased ice discharge from potentially unstable regions.

Atmospheric temperatures over the GrIS have increased significantly since the early 1990s, increasing by 1.7°C between 1991 and 2006 (Hanna et al., 2008; Box et al., 2009), which appears to have coincided with widespread glacier retreat (Moon and Joughin, 2008). Northern Greenland experienced negative mass balance between 2006 and 2012 (Khan et al., 2015), primarily due to enhanced surface melting and runoff (van den Broeke et al., 2009). Carr et al. (2013a) fit increased in air temperatures at NW and NE Greenland meteorological stations between 1990 and 2010 with a linear trend, which coincides with the dominant pattern of retreat at outlet glaciers in the region (Figure 7). In particular, the northeast region of Greenland has experienced increased discharge and melting between 2003 and 2012, which has been correlated to atmospheric warming (Khan et al., 2014). This coincides with substantial retreat of both Nioghalvfjerdsfjorden and Zachariae Isstrøm. While the majority of glaciers have shown retreat, the response is clearly non-uniform. Several of the study glaciers showed large variability in their terminus positions and sometimes advance (Storstrømmen, Kofoed-Hansen Bræ, and L. Bistrup Bræ; Box and Decker, 2011). Others showed a velocity increase of a factor of 10 (e.g., Harald Moltke Bræ and Hagen Bræ). Some of this behavior (e.g., periods of order of magnitude increased velocities accompanied by glacier advance) may be attributed to internal surging dynamics (Section Glacier Surging). Thus, whilst it is likely that increased air temperatures in northern Greenland have influenced glacier retreat over the last two decades, there has not been a coherent response.

In northern Greenland, greater surface meltwater production due to increased air temperatures has been linked to inter-annual retreat at Humboldt Glacier (Carr et al., 2015). Here, hydrofracture of crevasses a few kilometers inland of the glacier terminus may have caused weakening and promoted calving once the ice reached the terminus (Carr et al., 2015). At several glaciers, the presence of supraglacial lakes has also been noted (Humboldt, Ryder, Nioghalvfjerdsfjorden; Joughin et al., 1996b; Thomsen et al., 1997; Carr et al., 2015), and they are likely to be present on other outlet glaciers across northern Greenland. These lakes may enhance rates of calving through hydrofracture (e.g., Sohn et al., 1998; van der Veen, 1998; Carr et al., 2015), and the role of supraglacial lakes in northern Greenland, particularly across the NEGIS, could become increasingly important in the future (Ignéczi et al., 2016).

In other areas of the ice sheet it was initially thought that increased meltwater inputs led to seasonal-scale velocity increases (Zwally et al., 2002; Pimentel and Flowers, 2010). However, more recent work has linked increased meltwater production to a net annual slowdown in velocity, due to the drainage systems capacity to adjust and more efficiently drain adjacent high pressure areas via larger subglacial channels (Sole et al., 2013; Tedstone et al., 2015). Numerical modeling results suggest that the influence of meltwater inputs on seasonal velocity variations at Peterman Glacier is substantial (Nick et al., 2012), but little is known about this effect elsewhere in northern Greenland. To date, the potential impact of supraglacial lakes on northern Greenland outlet glaciers and their floating ice tongues has not been assessed. Based on observations from Antarctic ice shelves (Banwell et al., 2013), supraglacial lake drainages may play a role in calving events from large floating ice tongues by fracturing and weakening the ice. Further work to measure the occurrence, volume, and timing of supraglacial lake drainages is required due to the abundance of floating ice tongues in northern Greenland.

Alongside the role of surface meltwater induced changes discussed above, rates of submarine melt, primarily along the base of floating ice tongues, is likely to be an important control on glacier dynamics in northern Greenland. Submarine melt is likely to depend on both ocean temperature trends and topographic controls, whereby fjord configuration and depth control the access of sub-surface waters to glacier fronts. Submarine melt may be further enhanced by submarine meltwater plumes discharged at the grounding line (Motyka et al., 2003; Jenkins, 2011), where the more buoyant freshwater discharge promotes the circulation of deep warm water toward the grounding line (Motyka et al., 2003, 2011). Subglacial discharge is considered another primary control on submarine melt rates (Jenkins, 2011; Xu et al., 2012; Motyka et al., 2013; Sciascia et al., 2013), which could be strongly influenced by the amount of meltwater produced at the glacier surface and thus ultimately forced by atmospheric temperature changes.

Across the study region, ocean temperatures have been identified as a key control on outlet glacier behavior and ice tongue disintegration. In contrast to other areas of the GrIS, ice loss is thought be dominated by submarine melting on large floating ice tongues (Reeh et al., 1999; Rignot et al., 2001). At Nioghalvfjerdsfjorden and Zachariae Isstrøm, recent retreat rates may be due to high rates of submarine melt (~6 to 8 m a⁻¹; Rignot et al., 1997). Similarly, at Petermann and Tracy Glaciers, the intrusion of warm ocean water beneath floating ice tongues could have contributed to high rates of submarine melting, reaching up to 25 m a⁻¹ beneath Petermann Glacier (Rignot and Steffen, 2008), and subsequent ice tongue disintegration (Johnson et al., 2011; Johannessen et al., 2013; Porter et al., 2014). At Petermann Glacier, it was hypothesized that warmer ocean waters may have been a precursor to the large calving event in 2010 (Johannessen et al., 2013), where basal channels beneath the floating ice tongue underwent the greatest thinning (Münchow et al., 2014). Thus, Petermann Glacier's floating ice tongue is considered vulnerable to the temperature of relatively warm subsurface water entering the fjord (Johnson et al., 2011). We thus suggest the interaction between floating ice tongues and the ocean could have important implications for submarine-melt induced ice tongue collapse elsewhere in northern Greenland. However, limited in situ measurements of submarine melt rates, ocean temperatures (Thomsen et al., 1997), fjord circulation, and meltwater plumes means the extent of this process across northern Greenland remains unknown.

Forecasts suggest that ocean temperatures and submarine melt rates are likely to increase with future climate warming (e.g., Collins et al., 2013). As submarine melt rates of between 6 and 25 m a⁻¹ dominates mass loss at several northern Greenland outlet glaciers (e.g., Petermann, Nioghalvfjerdsfjorden, and Zachariae Isstrøm; Rignot et al., 2001), and has the potential to greatly influence glacier stability, there is need for improved fjord temperature data to better estimate submarine melt rates and the role of subglacial meltwater plumes (Nick et al., 2012). Systematic measurements of surface mass balance vs. submarine melting and calving are also required to better understand the importance of these processes of mass loss for many glaciers in northern Greenland. Northern Greenland glaciers are likely to be particularly vulnerable to ocean warming, due to the presence of extensive floating ice tongues, with large surface areas susceptible to submarine-melt induced collapse.

Fjord width has been shown to have a strong influence on glacier dynamics (e.g., Jamieson et al., 2012; Enderlin et al., 2013; Carr et al., 2015). Fjord width variations can influence the stability of marine-terminating outlet glacier front positions by either promoting equilibrium or advance in a narrowing fjord, or rapid retreat in a widening fjord (e.g., Benn et al., 2007; Carr et al., 2014). At Petermann Glacier, the narrow fjord is thought to hinder the movement of icebergs away from the glacier front, which may facilitate ice mélange formation, which, also referred to as sikussak, is defined as a mixture of calved icebergs and sea ice. This mélange may “choke” the fjord with icebergs, which could exert resistive back-stress on the glacier tongue (Johnson et al., 2011). This ice mélange has been identified as a key control on iceberg calving rates elsewhere in Greenland (Amundson et al., 2010; Cassotto et al., 2015). Narrow fjords may also result in more ice contact with the fjord walls and, consequently, greater lateral drag exerted on the glacier sides (Raymond, 1996). These processes may apply to other glaciers in the region, several of which also terminate in narrow fjords (e.g., Ryder, Steensby and Hagen Bræ), and there is large variability in fjord geometries across northern Greenland, ranging from the wide Humboldt Glacier to the narrow sinuous fjord at Steensby Glacier. Thus, contrasting fjord widths between glaciers in northern Greenland could contribute to their varying response to external drivers.

Changes in the floating ice tongues in front of several outlet glaciers in northern Greenland are another glacier-specific factor which could have influenced past glacier dynamics in northern Greenland. Table 1 highlights glaciers in the region which still have floating ice tongues and those where they have been lost. Changes in buttressing forces provided by floating ice tongues can influence glacier velocities (Howat et al., 2007; Nick et al., 2009, 2012). Floating ice shelf collapse led to increased glacier velocities at the Larsen B ice shelf in Antarctica (Scambos, 2004) and there is potential for this process to occur at floating ice tongue terminating outlet glaciers in northern Greenland. However, little increase in velocities at Petermann Glacier (Nick et al., 2012; Johannessen et al., 2013) and C. H. Ostenfeld Glacier (Joughin et al., 2010a) were found following ice tongue collapse, suggesting this may not be the case. This is of key consideration in the future, as further ice tongue retreat at northern Greenland outlet glaciers could substantially increase ice velocities, although modulated by ice shelf and fjord specific characteristics.

In contrast to most of the rest of the ice sheet (Bamber et al., 2013; Morlighem et al., 2014), a large proportion of northern Greenland rests below sea level and is characterized by deep fjords beneath outlet glaciers (Figure 2B). The regions with the greatest areas below sea level are at Nioghalvfjerdsfjorden and Zachariae Isstrøm, Humboldt, and Petermann Glaciers (Figure 2B). At Nioghalvfjerdsfjorden and Zachariae, the basal trough extends ~130 km to the interior of the ice sheet and reaches up to 550 m below sea level (Figure 2B). Evidence from Humboldt Glacier also shows that basal topography can have a major impact on glacier retreat and ice velocities (Carr et al., 2015). Here, retreat and ice velocities are an order of magnitude greater in the northern sector, which is underlain by a deep basal trough (up to 475 m deep) and an inland-sloping bed. The southern section, is comparatively shallow (~220 m deep) and slopes upwards inland (Rignot et al., 2001; Carr et al., 2015). Petermann Glacier also has a deep channel that extends to the interior of the ice sheet (Bamber et al., 2013; Morlighem et al., 2014) which could have a substantial impact on ice dynamics.

At some of the smaller glaciers in northern Greenland, basal topography has also been identified as a potentially important control on dynamics. In particular, contrasting basal topography may have been responsible for the differing rates of retreat at Heilprin and Tracy Glaciers, where a deeper bed beneath Tracy Glacier allows a greater ice area to be subject to warmer ocean waters and associated increased submarine melt rates (Porter et al., 2014). Similarly, at Hagen Bræ and Academy Glaciers, deep basal troughs could be susceptible to ocean warming and linked to glacier instability (Morlighem et al., 2014). However, an alternative explanation for increased flow velocities at Hagen Bræ, may be the loss resistance as the glacier retreated from being in contact with an island pinning point (Joughin et al., 2010a). Basal topography is also important at Ryder Glacier, through its impact on water storage and short-term velocity variations (Joughin et al., 1996b; Abdalati et al., 2001). Thus, basal topography is likely to be an important control on observed glacier retreat and could have important implications for future instability in northern Greenland.

Of further consideration is the nature of the bed, which can have a key influence on ice-sheet dynamics. In particular, beneath the NEGIS, a weak, deforming bed has been suggested responsible for its streaming flow (Joughin et al., 2001; Layberry and Bamber, 2001), similar to that observed at ice streams in Antarctica (e.g., Bindschadler et al., 2001). More recent work supports this hypothesis that water-saturated till contributes to the flow speed on the NEGIS (Christianson et al., 2014). However, this, alongside the effect of subglacial geology, remain under-studied in Greenland in contrast to Antarctica (Walter et al., 2014), and deserves further research.

Differences in bed topography, fjord geometry, and the presence of floating ice tongues at outlet glaciers across northern Greenland are likely to partly explain the varying responses in glacier dynamics observed. However, little examination of glacier specific factors on outlet glacier retreat have been conducted in northern Greenland and we therefore identify this as a key area for future research.