Destabilisation of an Arctic ice cap triggered by a hydro-thermodynamic feedback to summer-melt

Introduction Conclusions References


Introduction
Glacier mass loss constitutes the largest contributor to global mean sea-level rise (SLR), followed by ocean thermal expansion (Church et al., 2011).Over the last two decades, not only the glaciers and ice caps (Kaser et al., 2006), but also the ice sheets have lost mass at accelerating rates (AMAP, 2011;Shepherd et al., 2012).These ice mass changes coincide with atmospheric warming that causes record summer temperatures and glacier melt in the Arctic (Gardner et al., 2011) and across the Greenland ice sheet (Tedesco et al., 2013).Recently, the West Antarctic ice sheet has been Figures Back Close Full reported as one of the world's fastest warming regions (Bromwich et al., 2013).In addition to mass loss by melting, glacier dynamics have the potential to significantly amplify glacier response to climate change by altering the ice discharge to the ocean.Outlet-glacier, respectively ice-stream acceleration is generally attributed to hydraulic lubrication and oceanic-warming induced destabilization of floating ice tongues in the case of the Greenland ice sheet and reduced buttressing by thinning or loss of ice shelves in the case of the West Antarctic ice sheet (WAIS) (Shepherd et al., 2012).Recent observations of sustained ice flow and mass loss from Northeast Greenland and West Antarctica are also reported as evidence of a marine ice-sheet instability (Khan et al., 2014;Mouginot et al., 2014).The phenomenon, also known as tidewater-glacier instability, refers to glacier speed-up due to terminus retreat into deeper water (Meier and Post, 1987;Pfeffer, 2007).
To date, glacier-dynamic feedback processes remain poorly constrained and are therefore not yet incorporated in global projections of future glacier wastage (Stocker et al., 2013).SLR projections released with the IPCC Fifth Assessment Report range from 0.26 to 0.82 m over the 21st century (Stocker et al., 2013).The projections include ice discharge from the ice sheets into the ocean, however, based on current discharge rates and not accounting for a future dynamic response to climate change.The IPCC acknowledges, but is not able to quantify the probability of significantly higher SLR, such as associated with disintegration of marine-based sectors of the Antarctic ice sheet (Stocker et al., 2013).Rapid marine ice-sheet disintegration is evident from geological records both in the Northern and Southern Hemisphere and typically associated with air temperatures similar or warmer than those predicted for the end of the 21st century (Cook et al., 2013;Deschamps et al., 2012).Ice-rafted debris distributed across the North Atlantic ocean floor provide evidence of substantial calving associated with the rapid disintegration of the Laurentide ice sheet, so-called Heinrich events (Bond et al., 1992).Paleoclimatic records suggest rates of sea-level change much larger than currently observed or projected for the 21st century, e.g.SLR per century (35-50 mm a −1 ) at the end of the last glacial maximum, ∼ 14.5 ka ago (Deschamps et al., 2012).
Variations in glacier flow and hence ice discharge, encompass wide ranges of timescales and magnitudes, from diurnal velocity fluctuations to century-scale surge-type behaviour, and are mainly attributed to changes in basal drag (Clarke, 1987).Seasonal velocity variations are externally controlled by surface-melt induced acceleration through basal lubrication (Bindschadler, 1983).The rate of water supply to the glacier bed and the capacity of the subglacial drainage system to accommodate increased discharge regulate the basal water pressure and therewith basal drag (Schoof, 2010).Increased meltwater supply may thus accelerate, but also slow down ice flow, if it facilitates the establishment of a hydraulically efficient drainage system (Schoof, 2010).Several recent studies therefore suggest that hydraulic lubrication alone has a more limited effect on the future net mass balance of the Greenland ice sheet (Nick et al., 2013;Shannon et al., 2013) than previously anticipated (Zwally et al., 2002).However, these assessments do not account for changes in the extent of the basal area subjected to hydraulic lubrication (Nick et al., 2013) or the associated calving loss (Shannon et al., 2013).The effect of surface meltwater on the thermal structure of glaciers has only recently been considered in modelling studies (Phillips et al., 2010).The process of latent heat release and direct heat transfer from water in glaciers was termed cryo-hydrological warming (Phillips et al., 2010).Cryo-hydrological warming has the potential to change englacial and basal temperatures within years, whereas changes in the basal thermal regime by heat conduction would require decades to centuries (Phillips et al., 2010).Present outlet glaciers may accelerate significantly, if meltwater drainage to the bed would spread inland and thaw areas of the bed that are currently frozen (Alley et al., 2008).We propose that cryo-hydrological warming may have a drastic effect on glacier dynamics by weakening cold-based marginal ice regions that obstruct outflow of inland ice regions with basal temperatures already at or near the pressure melting point.Figures

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Full Cyclic surge-type behaviour is an extreme example of variations in glacier flow and characterized by long quiescent phases (decades to centuries) with slow flow velocities followed by short-lived active phases (months to years) with orders-of-magnitude increases in flow velocities (Raymond, 1987).During a surge, mass is transferred from an upper reservoir area to a lower receiving area, often associated with kilometre-scale advances of the terminus and greatly enhanced calving flux, in the case of tidewater glaciers.A thermally controlled soft-bed surge mechanism has been proposed for the polythermal glaciers of Svalbard (Hamilton and Dowdeswell, 1996).When thawed, unconsolidated, water-saturated sediments lack shear strength and have the potential to efficiently destabilize the overlying ice (Tulaczyk et al., 2000).Surge-type behaviour is generally considered independent of climatic forcing.Internal mechanisms may be sufficient to drive cyclic instabilities (MacAyeal, 1993).Changes in glacier geometry during the quiescent phase provoke corresponding changes in driving stress and basal thermal regime (Clarke, 1976).Nevertheless, the climatic mass balance influences the build-up of the reservoir area and therewith the surge periodicity.In addition, external factors drive feedback processes that lead to amplified dynamic response, pushing the glacier towards an instability threshold.
Here we present observations of multiannual ice-stream acceleration of a large Arctic ice cap, culminating in a surge and drastically enhanced ice discharge.Our observations demonstrate strong links between surface-melt and ice-flow acceleration.We propose a hydro-thermodynamic feedback that efficiently weakens the basal drag in initially frozen areas and facilitates enhanced basal motion over an expanding temperate basal area.By identifying the hydro-thermodynamic feedback to summer melt as a trigger for the mobilization of initially slow-moving ice regions, our study contributes to the understanding of dynamic glacier wastage and disintegration mechanisms of the ice-sheets in a warming climate (Alley et al., 2008;Joughin and Alley, 2011) At ∼ 7800 km 2 , Austfonna is the largest ice cap in the Eurasian Arctic (Fig. 1a).The ice cap consists of a main dome with an ice thickness of up to 600 m that feeds a number of drainage basins, some of them known to have surged in the past (Lefauconnier and Hagen, 1991).Austfonna has a polythermal structure.In the ice cap's interior basal temperatures are likely at pressure melting point (Zagorodnov et al., 1989), while the thinner ice-cap margins are cold and frozen to the ground, except for distinct fast flowing outlets that are dominated by fast basal motion and intersperse the generally slowly deforming ice cap (Dowdeswell et al., 1999).The south-eastern basins are to a large extent grounded below sea level, and form a continuous, non-floating calving front towards the Barents Sea (Dowdeswell et al., 2008).The regions grounded below sea level are to some extent underlain by marine sediments (Dowdeswell et al., 2008).Over 2002-2008, the climatic mass balance of Austfonna was close to zero (Moholdt et al., 2010a).Yet, the ice cap was losing mass due to calving and retreat of the marine ice margin that accounted for 2.5 Gt a −1 or 33 ± 5 % of the total ablation (Dowdeswell et al., 2008).Basin-3 (∼ 1200 km 2 ) is the largest drainage basin of Austfonna and topographically constrained by a subglacial valley that extends from south of the main dome eastwards towards the Barents Sea (Fig. 1b).About one third of the ice base is grounded below sea level down to a maximum depth of ∼ 155 m within an overdeepening in the central terminus region.A previous velocity map based on interferometric SAR data acquired in the mid-1990s revealed an ice stream in the northern lower reaches of Basin-3 with flow velocities ≤ 200 m a −1 (Dowdeswell et al., 2008).The ice stream was topographically constrained by a subglacial mountain to the north (Isdomen) and nearstagnant ice, likely frozen to its bed, to the south.The absence of surface lineation (e.g.crevasses) as late as 1991, identified the ice stream of Basin-3 as a recent feature (Dowdeswell et al., 1999) , 2008).Basin-3 is known to have surged some years prior to 1870 (Lefauconnier and Hagen, 1991), and the advancing terminus created a pronounced surge lobe nearly 25 km in width.The terminal moraine associated with the last surge lies ∼ 8 km from the present-day position of the calving front (Robinson and Dowdeswell, 2011).
Based on lobe volume and accumulation rate, the duration of the quiescent phase has been estimated to last 200-500 years (Dowdeswell et al., 1991).

GPS velocity timeseries
In spring 2008, five GPS receivers were deployed along the mid-1990's central flowline, 5 to 16 km upglacier from the calving front (Fig. 1; Dunse et al., 2012).We used GPS single-frequency code receivers (L1 band).Geographical positions were logged at hourly intervals, every third hour for instruments installed after May 2011, at an accuracy typically better than 2 m (den Ouden et al., 2010).Filtering in the time domain was applied to reduce random errors, i.e. a 7-day running mean was applied to the daily mean position, velocities were computed, and finally, the velocity was smoothed by applying another 7-day running mean (Dunse et al., 2012).

Velocity maps from synthetic aperture radar
Velocity maps of Basin-3, Austfonna, were produced from 20 2 m resolution TerraSAR-X (TSX) scenes acquired between April 2012 and May 2013, and provided by the German Aerospace Center (DLR).Displacement fields were derived by using cross correlation between two consecutive acquisitions (Strozzi et al., 2002) and geocoded using a DEM of Austfonna (Moholdt and Kääb, 2012).was derived by the ice flux through a fixed fluxgate upglacier of the calving front and the area change downglacier of that fluxgate, multiplied by an average ice thickness.

Additional data
To approximate timing and magnitude of surface-melt periods, cumulative positivedegree days (PDD) were computed from the temperature record of an automatic weather station.The station was operated since April 2004 on the western flank of Austfonna at 22 • 25 12 E, 79 • 43 48 N and 370 m a.s.l.(Schuler et al., 2014).Surface crevasses were identified and mapped using ground-penetrating radar (Ramac GPR; Målå Geoscience) at a centre frequency of 800 MHZ, providing structural information from the glacier surface down to a depth of ∼ 12 m (Dunse et al., 2009, Appendix C).
GPR profiles were annually repeated since 2004 and geolocated using a kinematic Global Navigation Satellite System (GNSS; GPS and GLONASS).

Multi-annual ice-flow acceleration
In spring 2008, five GPS receivers were deployed along the mid-1990's central flowline, 5 to 16 km upglacier from the calving front (Fig. 1; Dunse et al., 2012).Winter velocities observed in May 2008, were significantly higher than in the mid-1990s (Fig. 2).The GPS time series also reveal considerable overall acceleration, occurring in pronounced steps, each of which coincides with the summer melt period, as indicated by the temperature record of an automatic weather station (Fig. 2; Schuler et al., 2014).High irreversible, resulting in elevated winter velocities (Appendix D, Fig. 8).In autumn 2012, strong acceleration continued although surface melt had ceased.

Mobilization of stagnant ice regions
The drastic acceleration during autumn 2012 coincided with the expansion of the fastflowing region across the entire basin, as revealed by a time series of ice-surface velocity maps based on intensity tracking of repeat-pass TerraSAR-X satellite radar images (Fig. 3; Appendix B).Ice-stream velocities in April 2012 were up to 3 m d −1 , one order of magnitude larger than in the mid 1990's (Fig. 3a).In April 2012, the ice stream extended further inland and had widened southward to a width of ∼ 6-8 km, as compared to ∼ 5-6 km in the mid-1990s.Consequently, the GPS receivers were located ∼ 1 km north from the central flowline and did not capture fastest flow velocities.In contrast to the mid-1990s, the south-eastern corner of Basin-3 displayed also fast motion, at velocities of up to 1 m d Consequently, the slow moving ice region in-between had decreased in size (Fig. 3b) and disappeared by October 2012, when ice flow escalated into a surge comprising the entire width of the basin, reaching velocities >10 m d −1 (Fig. 3c).Velocities increased further until January 2013, reaching a maximum of 20 m d −1 (Fig. 3d; Table 2).Between January and May 2013 the maximum velocities decreased to 15.2 m d −1 , while the upglacier regions continued to accelerate (Fig. 5b).By the end of 2013, fast flow of Basin-3 continued (Table 2).Introduction

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Calving flux
The TSX data allowed calculation of the calving flux components, i.e. (i) the ice flux through a fixed fluxgate near the calving front, and (ii) the mass change of the terminus downglacier of that fluxgate, accounting for front position changes (Fig. 4; Appendix B).The observed ice flux peaked at a rate of 13.0 ± 4.2 Gt a −1 in December 2012/January 2013, after which it decreased slightly (Fig. 4a).Prior to October 2012, the position of the entire calving front of Basin-3 was remarkably stable (slight retreat; Fig. 4b), indicating that the entire ice flux was balanced by iceberg calving.After November 2012, the southern and central parts of the front, advanced by >1 km, reducing ice mass loss through calving by 61 %, as opposed to a stable front position.
Direct conversion of calving mass loss to SLR contribution is only meaningful for a static calving front.Glacier surges are typically accompanied by significant terminus advances.An advancing terminus reduces the mass loss from the glacier, however, the submerged part of the terminus replaces sea water instantaneously, causing an instantaneous sea-level rise.We therefore distinguish between a glacier-mass balance and a sea-level perspective on the calving flux (Table 1; Fig. 4c).The activation of a spatial confined ice stream in the early 1990's (phase 1) can be explained by internal mechanisms, e.g.gradual changes in basal thermal regime and driving stress (Dunse et al., 2011), following the geometric build-up of the glacier over ∼ 120 years of quiescence.The multi-annual, stepwise acceleration in the prelude of the present surge (phase 2, observed by GPS since 2008) coincides with successive annual summer speed-ups (Fig. 2).Short-lived acceleration during the melt season is consistent with current understanding of hydraulic lubrication (Schoof, 2010).In contrast, the multi-annual acceleration of background velocities (Fig. 8) cannot be explained by this mechanism and suggests a fundamental change in dynamics.Enhanced post-summer velocities sustained throughout the winter can be explained by successive activation of previously stagnant ice regions during the previous summer melt period.Mobilization of increased ice volumes within the reservoir area lead to increased ice-stream velocities and discharge.A widening of the ice stream itself, allows for higher centre-line velocities due to the increased distance from the lateral shear margin (Joughin and Alley, 2011).
Mobilization of the reservoir area is evident from the surface-crevasse formation within the upper accumulation area of Basin-3.Ground-penetrating radar (GPR) surveys reveal first occurrence of surface crevasses from ∼ 2004 onwards (Appendix C; drag, which in turn regulates ice stream velocities (Joughin and Alley, 2011).Cryohydrologic warming and increased frictional heating both act to enhance this positive feedback, weakening the flow resistance exerted by the lateral shear margins and initially cold-based ice patches acting as "sticky spots".The interplay between cryohydrologic warming and the emergence of basal hydraulic lubrication over an expanding area of the ice base constitutes a hydro-thermodynamic feedback.Unconsolidated subglacial sediments underneath the marine-based ice of Basin-3 would act to further enhance this feedback, and when thawed, favour rapid destabilization (Hamilton and Dowdeswell, 1996;Tulaczyk et al., 2000).
A surge comprising the full width of Basin-3 and the subsequent advance of the terminus followed the mobilization of the remaining slow moving regions, notably including the calving front (phase 3).The abrupt onset of fast flow suggests sheer tearing of the glacier from its bed, after the remaining sticky spots have been sufficiently weakened by the hydro-thermodynamic feedback outlined above.While fast flow continued by the end of 2013, it is expected to slow down or come to a halt within a few years.Massive ice redistribution from the reservoir into the receiving area and towards the calving front efficiently lowers the driving stress.Eventually, temperate basal conditions underneath the dynamically thinned and decelerating terminus are expected to be no longer maintained, at which time the base of the ice will start to refreeze to its bed (Dunse et al., 2011).
To what extent can present dynamics of an Arctic ice cap like Austfonna shed light on the future dynamic behaviour of ice sheets?Analogue to the ice sheets, Austfonna consist of slow moving inland ice interspersed by faster flowing outlet glaciers and ice streams that deliver inland ice towards the calving front.Drainage basins in their quiescent phase are characterized by margins frozen to their bed (Dowdeswell et al., 1999).Similarly, the coastal margins of the Antarctic ice sheet contain large regions of cold-based ice (Pattyn, 2010) that may currently prohibit efficient drainage of warmbased interior ice towards the ocean -e.g. in the Wilkes Basin, East Antarctica, that is known to have been dynamically active during the Pliocene warmth (Cook et al., Introduction Conclusions References Tables Figures

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Full 2013).In southwest Greenland, further acceleration of Jacobshavn Isbrae followed the 2012 record summer melt season (Joughin et al., 2014).Although the authors attribute the acceleration to terminus retreat into a bedrock depression, the occurrence of pronounced summer speed up at higher elevations may also reflect inland migration of surface melt and the effect of associated hydraulic feedbacks on glacier dynamics (Alley et al., 2008;Meierbachtol et al., 2013;Phillips et al., 2013).Similarly, the hydro-thermodynamic feedback may play a role in the sustained mass loss recently reported from northeast Greenland and mainly attributed to sea-ice decline due to regional waming (Khan et al., 2014).

Summary and conclusion
Our Given continued climatic warming, we hypothesize that the hydro-thermodynamic feedback to surface melt will gain significance for the destabilisation of Polar land ice, including the ice sheets.In light of recent record melt and rising ELA of the Greenland Ice Sheet, the proposed mechanism has the potential to lead to a long-term enhancement of outlet-glacier discharge and calving loss.Our expectation contrasts recent studies that indicate limited effects of surface-melt induced acceleration on the future net-mass balance of the Greenland ice Sheet (Nick et al., 2013;Shannon et al., 2013).Surface melt in Antarctica is presently mainly constrained to the ice shelves (Comiso, 2000).Given strong continued warming, surface melt will increasingly occur over coastal areas of Antarctica, making the grounded ice-sheet margins vulnerable to the hydro-thermodynamic feedback.Our study of the Austfonna ice cap highlights the importance of dynamic glacier wastage for the mass balance of Polar land ice.Current model projections of future SLR (Stocker et al., 2013) still do not account for the dynamic response of glaciers to continued global warming and might need to be revised upward after incorporating mechanisms such as the hydro-thermodynamic feedback.
Appendix A: TerraSAR-X velocity maps 20 TerraSAR-X (TSX) satellite synthetic aperture radar scenes of ∼ 2 m ground resolution (TSX Stripmap mode) were acquired between April 2012 and December 2013 (Table 2), providing information on the spatial evolution of the surge (Fig. 3).For each pair of consecutive acquisitions, displacements were determined using cross-correlation of the intensity images (Strozzi et al., 2002) and 19 velocity maps produced by accounting for the respective repeat-pass period (Table 2).The size of the matching-window was adjusted according to the expected maximum displacements during the repeat pass cycle, 300 × 344 pixels in range and azimuth direction in case of 11 days or 599 × 688 pixels for 22 days or longer.Displacements were measured in discrete steps of 50 × 57 pixels in range and azimuth to achieve a resolution of ∼ 100 × 100 m.Velocity maps were calculated from displacement maps by accounting for the time interval between Introduction

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Full the two underlying TSX images (t 1 to t 19 , Table 2) and geocoded using a DEM of Austfonna (Moholdt and Kääb, 2012).Velocities larger than a carefully estimated maximum (Table 2) were classified as mismatches and removed.To estimate the ice flux (Appendix B), obvious erroneous velocities remaining in the vicinity of a fixed fluxgate were manually removed and the maps interpolated using inverse distance weighting to provide continuous velocity profiles along the fluxgate (Fig. 5a).A comparison between TSX and GPS velocities for each GPS station and all repeat-pass periods, revealed that TSX underestimated local velocities at the GPS stations by 0.3 m d −1 (Fig. 6).
We explain this small bias by large horizontal velocity gradients within large matching windows, which could lead to the average velocities within a matching window being typically smaller than the strictly local GPS velocity.Consequently, more stable areas produce best matches.

Appendix B: Calving flux
The calving flux, q, of Basin-3 was calculated based on TSX velocity maps and changes in extend of the glacier by q = q fg +q t , with q fg being the ice flux through a fluxgate near the calving front, and q t the volume change of the terminus downglacier of that fluxgate due to advance or retreat of the calving front.Here, we defined a spatially fixed fluxgate approximately perpendicular to the ice flow, typically 1-3 km upglacier from the actual calving front (Fig. 5a), and where ice surface velocities could be inferred from all TSX-image pairs.The ice-flux can be written as where H fg = zs fg −zb fg is the ice thickness along the fluxgate, with zs fg and zb fg surface and bedrock elevation, respectively, w fg is the width of the fluxgate and v fg the velocity across the fluxgate.We assume plug flow, i.e. depth averaged velocities equal surface velocities, because at the high flow velocities observed, ice deformation is negligible compared to basal motion (Clarke, 1987).Position changes of the calving front are Introduction

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Full addressed by changes in areal extent of the glacier downstream from that fluxgate per repeat-pass period where H t = zs t − zb t is the ice thickness at the terminus in vicinity of the calving front, where zs t represents a typical height of the calving front and zb t is the mean bedrock elevation for the area encompassed by the observed maximum and minimum glacier extent.∆A t is the areal change of the terminus over the repeat-pass period ∆t between successive TSX acquisitions.The uncertainty in the calving flux estimate is associated with uncertainties of the input variables listed in Table 3.If the variables are derived from several independent components, their uncertainties can be summed by the root of the sum of squares (RSS) of the uncertainty of the components.

B1 Ice flux
To calculate the ice flux q fg,i for each repeat pass period t i (Fig. 4a), the ice flux through 617, 50 m-wide fluxgate segments was integrated, accounting for local ice thickness and velocity.Ice volume is converted to mass using an ice density ρ ice = 917 kg m −3 .
The cumulative ice flux over the period 19 April 2012 to 9 May 2013 is derived by summing up the ice flux over 19 successive TSX repeat-pass periods, weighted by their time duration (Fig. 4b), case of t 3 , the available data aligns with the previous velocity profile for t 2 , a factor x relation was determined and the gaps filled by piecewise polynomial interpolation so that v(t 3 ) = x•v(t 2 ).In case of t 6 , the velocity profile was constructed by using the mean of the previous and following repeat-pass cycle, v(t 6 ) = 1 2 • (v(t 5 ) + v(t 7 )) (dash-dotted lines in Fig. 5a).The accuracy of the resulting velocity profiles were evaluated by comparing the constructed velocity profiles and the available reliable matches, revealing a standard deviation of 0.21 and 0.25 m d −1 and a negligible mean offset of 0.03 and 0.04 m d −1 , in case of t 3 and t 6 , respectively.
The uncertainty in the ice-flux term, ∆q fg , is derived according to the law of propagation of uncertainty , omitting indexes i and n.

B2 Position changes of calving front
Similarly, the calving-flux term related to position changes of the calving front can be expressed by the volume change of the terminus in vicinity of the calving front, q t : The calving front outlines of Basin-3 were digitized from geocoded TSX intensity images.Due to layover and shadow effects along the calving front a systematic error in the order of twice the height of the calving front of 30 m is assumed (Moholdt and Kääb, 2012).The digitizing accuracy is conservatively estimated to 4 pixels, i.e. ±8 m.Introduction

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Full Analog to the ice-flux term, the uncertainty in the terminus-change term, ∆q t , is determined by the law of propagation of uncertainty: For each repeat-pass period σ(q t ) is determined by , omitting index i .

Appendix C: Crevasse formation
Ground-penetrating radar (GPR) profiling was performed on Austfonna on an annual basis since spring 2004 (Dunse et al., 2009).The GPR was operated at a centre frequency of 800 MHZ and provides information over a depth range of about ∼ 12 m.The data allows for identification of individual surface crevasses.Survey profiles initially focused on the accumulation area.Since 2008, also the flowline along which GPS receivers are deployed, was repeatedly surveyed.In the accumulation area, crevasse formation took place between 2004 and 2007 (Fig. 7a and b) and 2008 and 2012 (Fig. 7c and d), for the western and eastern profile, respectively.The monitored flowline towards the terminus was already heavily crevassed at the time of the first GPR survey (Fig. 7c).

Appendix D: Background GPS velocities and summer speed-up
Our five-year GPS record from Basin-3 shows that the annual velocity minimum typically occurs in June, just prior to the onset of summer speed-up (Dunse et al., 2012).respectively (Fig. 8a and b).The strong acceleration results from flow velocities remaining above their pre-summer values after the annual summer speed-up.In particular, the reversible fraction of the summer speed-up decreases with time and diminishes in 2013, when the entire basin is surging (Fig. 8c).While the two upper GPS stations (B3 #4 and B3 #5) experience further drastic acceleration (factor 24 and 18 increase in June 2013 compared to June 2008, respectively), the lower GPS stations (B3 #1 and B3 #2) only show moderate velocity increases, because the locations are close to the lateral shear margins of the surging basin (Fig. 3c and d).Introduction

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. During the 1990s, the front retreated on average by 70 ± 10 m a −1 , accounting for two thirds of the calving flux of ∼ 0.4 Gt a −1 (Dowdeswell Discussion Paper | Discussion Paper | Discussion Paper | et al. Appendix A provides more detailed information on the TSX data and processing.Calving front outlines were digitized from geocoded backscatter images (Appendix B).The calving flux of Basin-3 Discussion Paper | Discussion Paper | Discussion Paper | sensitivity and short response time (days) of glacier dynamics to melt periods clearly suggest surface-melt triggered acceleration.The 2008 summer speed-up is followed by a gradual winter deceleration, returning to pre-summer velocities only before the onset of the 2009 summer speed-up (maximum not captured due to instrument-power loss).However, during subsequent years, the summer speed-ups were progressively Discussion Paper | Discussion Paper | Discussion Paper |

− 1 .
These two distinct fast-flow regions were completely separated by almost stagnant ice, notably including the calving front.Low ice velocities, < 0.1 m d −1 , indicate the absence of considerable basal motion, and suggest frozen-bed conditions in this region.In August 2012, i.e. at the end of the summer melt season, velocities had increased significantly, up to 6 m d −1 for the northern and 4 m d −1 for the southeastern fast-flow region, along with further lateral expansion of the fast flowing areas.
Discussion Paper | Discussion Paper | Discussion Paper | From 19 April 2012 to 9 May 2013, calving mass loss [yearly rate] from Basin-3 accounted for 4.4 ± 1.6 Gt a −1 [4.2 ± 1.6 Gt a −1 ], an order of magnitude increase compared to 1991-2008 (Dowdeswell et al., 2008), nearly tripling the calving loss from the entire Austfonna ice cap.The related sea-level rise contribution of 7.6 ± 2.7 Gt a −1 [7.2 ± 2.6 Gt a −1 ], is as large as the total glacier mass change from the entire Svalbard archipelago for the period 2003 to 2008, estimated to −6.6 ± 2.6 Gt a −1 (Moholdt et al., 2010b).5 Discussion of hydro-thermodynamic feedback mechanism The dynamic changes that have been observed at Basin-3 over the last two decades can be separated into three phases: (1) activation of a spatially confined ice stream in the early 1990's (Dowdeswell et al., 1999); (2) multi-annual acceleration from < 2008 Discussion Paper | Discussion Paper | Discussion Paper | to 2012, along with an expansion of the ice stream; (3) active surge phase following the destabilization of the entire terminus in autumn 2012.
Fig. 7).Crevasses are a manifestation of longitudinal extension and evidence of upglacier migration of the fast-flowing region.This is in line with the upglacier expansion of the ice stream revealed by TSX and the annual position change of a massbalance stake close to the ice divide (Fig. 1), subjected to strong acceleration from 11 m a −1 between 2004 and 2007 to 28, 43 and 114 m a −1 for the period 2010-2013.The first occurrence of crevasses signify the development of potential meltwater routes to the glacier bed, subjecting an increasing region to hydraulic basal lubrication and cryo-hydrologic warming.Similarly, meltwater reaching the bed beneath the heavily crevassed shear margin efficiently weakens the basal drag that balances the lateral Discussion Paper | Discussion Paper | Discussion Paper | observations of multiannual acceleration of an Arctic ice cap indicate successive enhancement of basal motion driven by meltwater supply and shed light on a destabilisation mechanisms of Polar land ice.We propose a hydro-thermodynamic feedback mechanism, triggered by surface melt reaching a growing fraction of the glacier bed.Intrusion of surface melt to the glacier bed provides an efficient heat source through cryo-hydrological warming and initiates hydraulic lubrication, weakening the resistance to basal motion, especially in initially cold-based ice regions.During the summer melt season, slow moving ice regions are successively mobilized through lateral and inland expansion of regions dominated by fast basal motion.The recent calving flux of Basin-3 is enormous and has strong implications for the mass balance of the ice cap and its contribution to sea-level rise.From 19 April 2012 to 9 May 2013, the calving flux [yearly rate] of Basin-3 amounted to 4.4 ± 1.6 Gt a −1 [4.2±1.6 Gt a −1 ], an order of magnitude increase compared to 1991-2008 (Dowdeswell et al., 2008).Accounting for the terminus advance, the related sea-level rise contribution of 7.6 ± 2.7 Gt a −1 [7.2 ± 2.6 Gt a −1 ] equals the total annual glacier-mass loss from the entire Svalbard archipelago for the period 2003 to 2008, estimated to 6.6 ± 2.6 Gt a −1 (Moholdt et al., 2010b).Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | n=1 H| fg,n • v| fg,i ,n • ρ ice To derive continuous velocity profiles along the fluxgates, the velocity maps were interpolated horizontally in the vicinity of the fluxgate where mismatches occurred.The velocity maps associated with long repeat pass periods, t 3 and t 6 , suffer from large gaps, only 51 and 46 % of the fluxgate segments hold sound velocity estimates.In Introduction Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Since 2008, the mean velocity in June has increased dramatically, with values in 2012 3 to 5 times higher than in 2008 at lowest (GPS B3 #1) and highest elevation (B3 #5Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Strozzi, T., Luckman, A., Murray, T., Wegmüller, U., and Werner, C. L.: Glacier motion estimation using SAR offset-tracking procedures, IEEE T. Geosci.
Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper |

Fig. 1 .
Fig. 1.Surface and bedrock topography of Austfonna/Basin-3.(a) Surface elevation contours at 50 m interval (solid black), overlain on a TerraSAR-X backscatter image (30 Apr 2012).The insert provides the ice cap's location within the Svalbard archipelago.Drainage basins are outlined in solid grey, solid green for Basin-3.Position of five GPS receivers and one stake on Basin-3, as well as the automatic weather station (AWS) are marked.Repeat-GPR profiling revealed crevasse formation in upper reaches of Basin-3 between 2004-2007 and 2008-2012 (Appendix D).(b) Bedrock contours (black, color-filled) are at 25 m intervals, with the bedrock sea-level contour highlighted in red and 50-m surface elevation contours superimposed (white).

Figure 1 .Fig. 2 .Figure 2 .Fig. 3 .
Figure 1.Surface and bedrock topography of Austfonna/Basin-3.(a) Surface elevation contours at 50 m interval (solid black), overlain on a TerraSAR-X backscatter image (30 April 2012).The insert provides the ice cap's location within the Svalbard archipelago.Drainage basins are outlined in solid grey, solid green for Basin-3.Position of five GPS receivers and one stake on Basin-3, as well as the automatic weather station (AWS) are marked.Repeat-GPR profiling revealed crevasse formation in upper reaches of Basin-3 between 2004-2007 and 2008-2012 (Appendix D).(b) Bedrock contours (black, color-filled) are at 25 m intervals, with the bedrock sea-level contour highlighted in red and 50 m surface elevation contours superimposed (white).

Figure 3 .Fig. 4 .
Figure 3. Surface velocity fields of Basin-3, Austfonna, derived from TerraSAR-X feature tracking; (a) April/May 2012; (b) August; (c) October; and (d) January/February 2013.Red circles represent mean position of GPS receivers over the particular repeat-pass period, fillcolor according to color-coding of receivers in Fig. 2. The red arrows indicate associated GPS velocity vectors.Glacier elevation contours plotted in grey at 100 m intervals, front position at time of repeat pass in orange.

Figure 4 .
Figure 4. Calving-flux from Basin-3, Austfonna, April 2012 to May 2013.Calving components are expressed in terms of the instantaneous (a) and cumulative mass change (b) and allocated to the effect on glacier mass balance and sea-level (c; see Sect.4.3).Whiskers (a, b) indicate uncertainty bounds calculated using propagation-of-uncertainty analysis and shaded areas (c) upper and lower bounds given maximum or minimum ice thickness.

Figure A1 .
Figure A1.Evolution of ice-flow velocities across a calving fluxgate (a) and along a central flowline of Basin-3 (b).Velocity profiles extracted from TerraSAR-X velocity maps April 2012 to May 2013 and median filtered to remove outliers.The insert shows the location of the profiles overlain onto the velocity map of t 14 .Fluxgate velocities are used to compute the calving ice flux.Velocities for long repeat passes t 3 and t 6 have been constructed as described in B2.

Figure A2 .Figure C1 .
Figure A2.Comparison of TSX and GPS velocities at GPS locations and for each repeatpass period.(a) Linear regression (solid black; R 2 = 0.94) predicts 94 % of the variance in TSX velocities (0.14 m d −1 ).(b) TSX velocities typically slightly underestimate local (sub-pixel scale) velocities measured at the GPS stations by 0.3 m d −1 .

Figure D1 .
Figure D1.GPS-velocity evolution along a central flowline of Basin-3; (a) annual velocity minimum defined as the mean value in June 2008 to 2013 and (b) normalized to mean-June velocity 2008; (c) reversibility of summer speed-up defined as the ratio of the velocity increase from pre-summer minimum (mean June) to maximum summer velocity and subsequent velocity decrease to pre-summer velocity of subsequent year.In 2009, summer maximum velocities were not captured due to GPS power loss.The triangles only indicate a minimum estimate for 2009-2010, based on elevated late-summer velocities measured after maintenance of GPS receivers in end of August.

Table 1 .
Estimate of calving flux components and total calving flux over the TSX observation period from 19 April 2012 to 9 May 2013.Q mb = Q fg − Q t 4.4 ± 1.6 4.2 ± 1.5 SLR perspective, Q sl = Q mb + Q tsd 7.6 ± 3.9 7.2 ± 2.6 Introduction

Table A1 .
TerraSAR-X acquisitions of Basin-3 and repeat-pass period and maximum velocities of inferred velocity maps.