Satellite remote sensing imagery

To analyze the changes to glacial ice in the study area, Landsat 7 Enhanced Thematic Mapper Plus (ETM+) and Terra ASTER visible and near infrared (VNIR) images were used to evaluate shelf area and other physical changes to the SNIS, RI and RG portions of the study area. The imagery data were projected to WGS UTM Grid 20S (ETM+) and 21S (ASTER) and the program QGIS (version 2.10.1 ‘Pisa’, 2015) was used to assess areas within our manually mapped SNIS or RI limits as well as the inland length of the RG. Stereo ASTER images from 2005 were also used to generate a digital elevation model (DEM) of RI (Table 2). For RI area change estimates, we chose to use ASTER imagery to avoid the need to extrapolate across the gaps in the ETM+ imagery that are due to the Scan Line Corrector failure (SLC-off) in 2003; however, this was not practical for the larger SNIS area. Finally, a long-term archive of MODIS imagery, available from the National Snow and Ice Data Center (NSIDC: https://nsidc.org/data/iceshelves_images/index_modis.html, Larsen B) (Scambos and others, Reference Scambos, Bohlander and Raup1996), was used to qualitatively evaluate the presence or absence of extensive fast ice in the adjacent embayments as well as indications of surface melt in the austral summer months particularly when higher resolution imagery (Table 2) was not available.

Grounding line (GL) and area estimates

We chose to use the GL positions derived from interferometric synthetic aperture radar (InSAR) in the late 1990s by Rack and Rott (Reference Rack and Rott2004), hereafter called the R&R-GL, to define the interface between the SNIS and RG and to determine which portions of the ICESat profiles were on and off the SNIS (Fig. 1). Note, the GL was used to define the western limit of the SNIS even though close inspection of recent imagery suggests it has moved inland at RG since it was defined with 1995–99 interferometry. The MOA2004 GL data, defined by breaks or inflections in ice slope (Haran and others, Reference Haran, Bohlander, Scambos, Painter and Fahnestock2005), do not differ significantly from the R&R GL where the SNIS meets the AP. For the small observed differences, it is not clear if they are due to the different techniques or changes in the ice over time (Depoorter and others, Reference Depoorter2013, Supplementary Information, Section 5). As in Haran and others (Reference Haran, Bohlander, Scambos, Painter and Fahnestock2005), visual inspection of satellite images was used to identify later GL positions for SNIS and RI.

The image data made it apparent that GL position had changed during the study period as had ice shelf area (Fig. 1; Table 2). To define the SNIS and RI areas, we used surface features observed in the imagery sequence to approximate the overall shelf remnant ice area, including all internal nunataks and to also distinguish it from adjacent grounded ice areas including the RI ice cap. Note, we do not include the ice shelf remnant west of Cape Fairweather and its protruding underwater ridge (~44 km² in 2002, lost in early 2004) defined by the GL adjacent to the SNIS area (Fig. 1).

The area values from this analysis (Table 2) have a combined uncertainty from a small amount of measurement error and a larger, but less defined, grounding zone position uncertainty. In the overall SNIS region, we estimate ~10 km² uncertainty caused by shelf extent/grounded ice extent discrimination. For RI alone, we compared extents digitized from two ASTER images taken 5 d apart in 2005 (10 and 15 November 2005), which showed a net area difference of <0.4 km². This test provides a sense of the measurement error and repeat assessments of single RI scenes produced similar or smaller area differences, but all required using an approximate GL position relative to the adjacent SNIS ice area. The use of ETM+ SLC-off imagery also increases area estimate uncertainties for SNIS as those images require extrapolation across narrow regions without visible data. Based on this, we expect areas to be uncertain at the few km² level for the ASTER–derived RI areas and ~10 km² for the ETM + estimates of SNIS.

Satellite laser altimetry

For all ICESat tracks, we selected repeat profiles with the least cloud obscuration (i.e. complete or essentially complete profiles) of the ice shelf surface and the adjacent ocean and grounded ice areas for further analysis (Table 3). After this initial filtering, the Release 633 data was de-tided, Geoid-Centroid bias corrected (Borsa and others, Reference Borsa, Moholdt, Fricker and Brunt2014), and then adjusted to a sea level datum. This was done using the average of the lowest 3 elevations in the nearest 2 km of the profile over the adjacent embayments (Kwok and others, Reference Kwok, Zwally and Yi2004; Scambos and others, Reference Scambos, Sergienko, Sargent, MacAyeal and Fastook2005). These data were then compared sequentially to assess thinning and margin retreat and anomalous and incomplete tracks were removed. As in Pritchard and others (Reference Pritchard2012), the CATS2008A tide model (Personal communication from L. Padman, 2014) could not be readily applied because of distinct uncertainties in grounded ice positions within this small area. A discussion of these issues for tide-corrections is found in Griggs and Bamber (Reference Griggs and Bamber2011).

Under clear sky and for grounded ice having low slope, ICESat measurements have an uncertainty of ~±0.14 m (Shuman and others, Reference Shuman2006). Later releases of ICESat data had a validated measurement uncertainty of 0.15 m (Schutz and others, Reference Schutz, DiMarzio, Luthcke, Hancock and Urban2011) although performance varied slightly over the mission. As such, RMS elevation differences of ~0.20 m could exist between exact repeat tracks over an unchanging grounded ice target. For floating ice shelf profiles, the components of elevation uncertainty are discussed in detail as being ~±0.275 m for a single track (Pritchard and others, Reference Pritchard2012, Supplementary Information, page 9). Based on this, an RMS of ~0.40 m might be expected for elevation profile differences from floating ice areas.

Elevation and thickness change estimates

For the SNIS remnant, we sought to evaluate elevation changes over the maximum time period possible (Table 3) within the ICESat mission. For the grounded ice profiles we used repeat profiles with small laser footprint separations, sometimes limiting the temporal separation of observations. However, because these offsets were usually less than the nominal ICESat laser footprint (~70 m, every ~170 m along track) (Schutz and others, Reference Schutz, Zwally, Shuman, Hancock and DiMarzio2005) we were able to minimize cross-track slope impacts on the derived elevation change estimates (Hofton and others, Reference Hofton, Luthcke and Blair2013).

SNIS elevation loss estimates

Figure 2 shows the selected repeat profile data for each track location with the maximum temporal separation. Cloud cover impacted some tracks (e.g. Fig. 2a). Cross-track separations of the Figure 2 profiles ranged from 14 m for Track 0256 to 111 m for Track 0367. However all profiles are in very low slope areas (the SNIS ice shelf surface) except where crossing grounded areas. We further reduced the impact of slope, and of individual measurement noise by taking the average elevations of each ICESat altimetry profile (e.g. the profile positions of the red lines in Fig. 2) and deriving regression lines through those average elevations (Fig. 3).

Fig. 2.

Elevation profiles and elevation loss rates are derived from the first and last available ICESat track across four parts of the Seal Nunataks ice shelf remnant. The first and last repeat track data show the overall changes to the shelf remnant including iceberg calving and elevation losses. The red lines indicate the magnitude and pattern of the loss rate across each profile. All plots have the same y-axis range but different x-axis ranges. Missing points are indicated by gaps in the plotted data. ICESat's orbit direction is indicated with the arrow and the degree of vertical exaggeration (VE) of the elevation data is shown in each plot. The black bar indicates 5 km in distance.

Fig. 3.

Plot of the mean elevation of each ICESat repeat profile crossing the remainder of the Seal Nunataks ice shelf remnant during 2003–09. The ICESat campaigns used in the study are labeled at the bottom of the plot. The regression lines indicate the elevation loss as a function of time and were used to calculate change rates in Table 4. The trend of Track 0256 losses is extrapolated to 2009 (dashed portion) on the basis of its regression line.

Although seasonal variations or densification of recent snowfall might impact the observed elevation change results in other regions, we believe that these are not significant issues here given the low annual snow accumulation for the southeastern nAP (van Wessem and others, Reference van Wessem2016), and strong seasonal surface melting and re-freezing of the annual snowfall. The linearity of our observed SNIS changes using ICESat campaigns, at periods out of phase with seasonality, suggest that seasonal variations must be small. Moreover, a six-season observational record on the SIIS and four-season record on Cape Disappointment from Automated Meteorology–Ice/Indigenous species–Geophysics Observation Systems, AMIGOS (Scambos and others, Reference Scambos, Ross, Haran, Bauer and Ainley2013) have never indicated a snowfall event of 10 cm snow height, and show that the seasonal elevation variation in the region is not more than 20 cm. AMIGOS-related field work by one of us (TAS) and an analysis of elevation and thickness data on SNIS, Larsen B and SIIS (Holland and others, Reference Holland2011) documents that the ‘firn’ in this area is ice. The repeated indication of surface melt ponding in summertime satellite imagery is a further indication of impermeable ice at the surface. Based on these lines of evidence, we do not apply a firn-compaction or seasonal elevation correction to these data.

SNIS thickness loss estimates

In contrast to the analysis in Pritchard and others (Reference Pritchard2012) and specifically for the SNIS (Personal communication from Pritchard, 2014), firn compaction is essentially nil in the area. Therefore, for the SNIS profiles we assumed all elevation change is directly related to a change in shelf thickness. Further, we assumed that the ice losses detailed in Figure 2 are from shelf areas in hydrostatic equilibrium away from GL positions.

To convert ice elevation changes (Fig. 3) to thickness changes, we modified Eqn (1) from Griggs and Bamber (Reference Griggs and Bamber2009) by eliminating the firn terms. Ice thickness changes at the floating ice positions profiled by ICESat were then determined using the relationship:

$$\Delta Z = \Delta e\; \times \;\; \rho_{\rm w} /(\rho_{\rm w} - \rho_i )$$

where ΔZ is change in ice thickness, Δe is elevation change from the regression lines in Figure 3, ρ _i is the estimated mean density of the ice shelf and ρ _w is the mean sub-ice-shelf water density. We used an average ice shelf density of 900 kg m⁻³ and a seawater density of 1028 kg m⁻³ (Personal communication from B. Huber, 2013) to convert observed elevation to estimated thickness change. This results in an ice shelf freeboard of 0.128 m for each meter of ice shelf thickness that allows derivation of evolving shelf thickness estimates from the data in Figure 3. For RI and RG profiles, the elevation loss is similarly assumed to be all ice without firn compaction although that approximation may not be the most appropriate for the highest elevation profiles in the study area (e.g. Track 0248 on RG, Fig. 1).

Grounded ice elevation changes

We differenced repeat profiles with small laser footprint separations for all the grounded ice areas with data on RI and RG. Because these offsets were less than the nominal ICESat laser footprint of ~70 m (Schutz and others, Reference Schutz, Zwally, Shuman, Hancock and DiMarzio2005), cross-track slope impacts on the derived elevation changes were minimal. The footprint separations averaged 38 m for RI's Track 0003 data, 41 m for the RG's Track 0248 data and 44 m for the Track 0256 data. We tried to correct the influence of the cross track slopes on RG using an external SPOT5 DEM (Korona and others, Reference Korona, Berthier, Bernard, Remy and Thouvenot2009) but the procedure was unsuccessful due to the high noise level in the DEM in this area with limited contrast in the imagery. However, the 2005 ASTER DEM for RI (Fig. 4) was successfully utilized as described in Shuman and others (2011, Section 2.2) to correct for cross-track slope variations in the available ICESat data. No Operation IceBridge airborne altimetry data are available over this area. Due to the lack of independent elevation data, profile elevation difference uncertainties are estimated from the standard deviation of the shot-to-shot elevation differences (σ values in Table 4).

Fig. 4.

Changes in the ice cover of Robertson Island (RI) over the study period. Panel a shows the position of Track 0003 across the island and also Track 1349 crossing Hertha Nunatak (He.) as well as the adjacent ice shelf remnant in 2012. Multiple ASTER images were used to define ice edge positions but only the 2001 and 2012 edges are shown here (see Table 2). A 2005 ASTER image pair (Table 2) generated the DEM used for the elevation contours. The dashed lines indicate interpreted ice shelf remnant limits and solid lines are used for island areas. Nearby features are labeled in the same manner as in Figure 1 (Table 1). Panel b shows the first and last Track 0003 elevation profiles as well as the derived elevation change rates. ICESat's direction is indicated with the black arrow and the degree of vertical exaggeration (VE) of the elevation data is shown. The black bar indicates 5 km in distance. Missing ICESat shots are indicated by dashed lines in the island's profile.

Table 4.

Summary of repeat-track elevation and estimated ice shelf changes

Loss values for the SNIS are derived from the regression lines shown in Figure 3. The losses for RI and RG are the mean elevation differences between the spatially-closest pair of ICESat profiles. As discussed in the text, RG is divided into north (N) and south (S) branches (Fig. 5a). For all the profiles, the elevation loss is also assumed to be without firn compaction. The standard deviations (σ) of the derived SNIS elevation losses are from first – last repeats (see Table 3) whereas the RG and RI σ values are from spatially closest profile pair. Because of RI's topography accentuates mass losses on the north side, we used data from the lower 200 m on the south side of the island (25 points) to derive the σ value.

SNIS ice shelf area changes

Over the 2001–13 time period, the SNIS ice edges have retreated considerably (Fig. 1; Table 2) (see also Ferrigno and others, Reference Ferrigno2006, Reference Ferrigno2008; Cook and others, Reference Cook, Fox, Vaughan and Ferrigno2005; Cook and Vaughan, Reference Cook and Vaughan2010). The SNIS area, not including RI (discussed below) but including embedded nunataks, reduced from 992.2 km² in 2002 to 743.1 km² in 2013, a net loss for this period is 249.1 km² with an estimated uncertainty of ~10 km². The losses on SNIS's north side have been previously indicated for the Larsen A by Cook and Vaughan (Reference Cook and Vaughan2010, their Table 3, Fig. 4.9) but the data in Table 2 specifically track the SNIS's area.

During mapping of the ice shelf area from the image series, the presence and progression of a series of rifts between Ocean and Castor nunataks (see Fig. 4a) was observed. These rifts do not hold water even when surface melt is evident nearby and lead to the release of icebergs from this area. The imagery also shows older undulations that frequently held melt water during the study period. Based on their orientation, these tight undulations must have formed by the convergence of ice from adjacent glaciers (Fig. 1 and Fig. 4a). We infer that since the new rifts cut across the older ones, the open rifts reflect new stresses on the ice shelf remnant.

Several of the SNIS ICESat profiles show the occurrence of ice edge retreat, but all indicate significant net elevation changes (Fig. 2). For example, ICESat Track 0367 data (Fig. 2b) shows the vertical scale of the changes near the widest part of the SNIS. This portion of SNIS is not within the area of embedded nunataks but is where successive icebergs calved during the study period as is suggested by the repeat altimetry profiles and Figure 1. MODIS imagery from 3 March 2000 shows the presence of the large rifts that produced these icebergs, indicating they were probably a response to loss of buttressing from the Larsen A ice shelf following the collapse in 1995 (Ferrigno and others, Reference Ferrigno2006; Cook and Vaughan, Reference Cook and Vaughan2010). The icebergs calved from the north side of the shelf remnant in 2004/05 (Shuman and others, Reference Shuman, Berthier and Scambos2011). The profiles for Tracks 0256 and 0375 show smaller iceberg losses. Other small calving events are indicated in Figure 1 such as the ice area lost between Castor Nunatak (Ca.) and RI. By combining imagery data (Table 2, MODIS) and all ICESat profiles (Table 3), the timing of these ice area losses of shelf ice can be constrained.

Floating ice elevation and thickness changes

The ICESat repeat altimetry allows comparisons of elevation changes for several parts of the SNIS remnant. Bathymetric data adjacent to the shelf areas crossed by the profiles (Lavoie and others, Reference Lavoie2015) show that the shelves are afloat at least at the edges, and distinct grounding lines in InSAR, MODIS and Landsat 7 image data far from the shelf edges show that the ICESat track areas are largely afloat (Rack and others, Reference Rack, Rott, Siegel and Skvarca1999; Haran and others, Reference Haran, Bohlander, Scambos, Painter and Fahnestock2005; Bindschadler and others, Reference Bindschadler2011).

Figure 2 shows the first and last altimetry profile for each of the four SNIS tracks as well as the derived loss rate across the remaining shelf remnant from those two tracks. Lateral offsets between the track pairs shown in Figure 2 ranged between 14 m (0256) and 111 m (0367) and averaged 55 m. Using a 5-shot smoothed, average along-track slope value from Track 0367 suggests the cross track distance at Track 0367 could cause at most ~0.065 m of elevation difference, or ~5% of the observed 1.43 m net elevation change. The shape of the loss rate data (red lines in Fig. 2) suggests slightly greater losses at the exposed ice fronts of the shelf remnant (Tracks 0367 and 1349) but the data are less diagnostic for Tracks 0256 and 0375.

The slope of the regression line for all ICESat repeat track data was used to calculate the elevation loss rates in Table 4 for each track location. The altimetry profile pairs shown in Figure 2 illustrate the pattern of the elevation change. Given the sparse temporal record of the altimetry profiles, it was not possible to assess any seasonal elevation variations but we believe those effects must be small. Single-profile elevation uncertainty, indicated by error bars in Figure 3, are based on Pritchard and others (Reference Pritchard2012, Supplementary Information, Section 1.4 (±0.275 m)). The similarity of the trends for the four different SNIS profiles increases confidence in these results.

Table 4 shows the elevation loss rates and derived ice shelf thinning rates for each SNIS track. Our calculated rates are about twice the values reported in Pritchard and others (Reference Pritchard2012; Personal communication from Pritchard, 2014) as discussed further below. However, our individual track analyses provide consistent results across the ice shelf remnant (Table 4, Fig. 1) and an overall estimated mean elevation loss of 0.29 m a⁻¹ and a mean thinning of 2.32 m a⁻¹.

Significant differences in firn compaction estimates from modeling results versus field observations seem to explain why our loss rates are markedly higher than earlier work. Modeling data for the Larsen B (Pritchard and others, Reference Pritchard2012, Supplementary Information, Table 1) suggests there is 0.16 m a⁻¹ elevation lowering on the nearby Larsen B remnant attributed to firn compaction. As noted, this is inconsistent with field observations of solid ice in the upper meter of the SIIS surface during field work (TAS, February 2010), analysis of the mean density of the entire shelf (Holland and others, Reference Holland2011) and the satellite image record of extensive melt ponding in the area since the mid-1980s (Scambos and others, Reference Scambos, Hulbe, Fahnestock, Domack, Leventer, Burnett, Bindschadler, Convey and Kirby2003).

Grounded ice elevation changes

Quantifying the elevation changes for RI was achieved by using a pair of ICESat repeat profiles with overlapping footprints (mean offset was 38 m) as well a DEM derived from a pair of ASTER images (Table 2) as discussed further below. During the same overall study period for the SNIS, repeated ASTER imagery showed that RI's ice margins retreated up to ~0.5 km, primarily on the northern margin (Fig. 4a and Table 2). The limited retreat observed in the imagery suggests that the majority of RI is covered in grounded ice.

The RI ice cap became separated from adjacent major ice shelves in the late 1990s (Ferrigno and others, Reference Ferrigno2008, Map I-2600-B) following major retreats of the Larsen A and B ice shelves. Our imagery analysis distinguished RI from the SNIS remnant by a subtle slope break (see the section  SNIS ice shelf area changes) and indicates the ice cap reduced in area from 166.3 km² in 2001 to 159.2 km² in 2012. The net loss is ~7 km² with an estimated uncertainty of ~2 km² because the ASTER imagery cannot exactly define the RI interface position with the SNIS. The small area change and absence of calving suggests that the RI ice cap is grounded.

Plotting the elevations and the derived differences of the first and last Track 0003 profiles across RI, 22 February 2004 (2B) and 14 March 2009 (2E) (Table 3), illustrates significant changes in the ice cap's topography (Fig. 4b). The difference data from these profiles also defines a variable loss rate across RI. This ranges from ~1 m a⁻¹ on the south side to <0.5 m a⁻¹ at the summit ridge to a maximum of nearly 4 m a⁻¹ in a topographic ‘bowl’ on the northeast-facing side of RI. Table 4 shows the mean elevation loss and loss rate derived from the two profiles as well as a larger uncertainty value (see the section  Satellite laser altimetry) derived from the standard deviations of the differences from a portion of the RI data.

In addition, all the Track 0003 profiles (Table 2) were differenced using a 2005 ASTER DEM to account for cross-track slope variations between the offset altimetry profiles (after Moholdt and others, Reference Moholdt, Nuth, Hagen and Kohler2010). Track spacing in the full set of the clear Track 003 profiles ranged up to 350 m. Net elevation change between profiles was assessed using a regression line through the mean elevation differences from the DEM for each repeat profile. The difference between those results and the mean elevation loss from the first-last (F-L) analysis shown in Figure 4b was 0.04 m a⁻¹, with the DEM-based result being smaller. This indicates that our results for small footprint-offset passes are consistent with methods that apply a cross-slope correction (e.g., Scambos and others, Reference Scambos2014). Note, for the analysis in this study, we filtered less of the ICESat data on RI than in the 2014 study. This produces a slightly more negative elevation change rate (−1.76 m a⁻¹, Table 4) than was reported in the earlier study (−1.54 m a⁻¹) (Scambos and others, Reference Scambos2014, supplement, Basin 26b).

Rogosh Glacier (RG), the last remaining tributary to the western end of SNIS is crossed in two places by ICESat track data (Tracks 0248 and 0256, Fig. 1 and Table 3) but the altimetry data are sparse due to clouds. Figure 5 shows the elevations of the closest pair of profiles for Tracks 0248 and 0256 both inland of the GL (Fig. 1). The derived loss rates are noisy in part due to the topography of RG. Altimetry data gaps, likely due to clouds, impacted the 2E data for 0248 (Fig. 5a) and the inland portion of Track 0256 (e.g. Fig. 2a) was not available due to clouds after November 2006 (Table 3).

Fig. 5.

ICESat repeat altimetry profiles and derived elevation loss rates for the Rogosh Glacier (RG) (see Fig. 1). Figure 5a shows the data where ICESat Track 0248 crosses the middle part of the branching RG nearly perpendicularly. Figure 5b shows the data where ICESat Track 0256 obliquely crosses the lower RG's rough topography. Dashed lines indicate approximately where the valley walls were located and these plots have different axis ranges. Missing altimetry data are indicated by gaps in the plotted data. ICESat's direction is indicated with the black arrow and the degree of vertical exaggeration (VE) of the elevation data is shown. The black bars indicate 1 km in distance.

The Track 0248 data reveal a difference in both topography and elevation loss rate across the upper part of this tributary glacier. These altimetry profiles cross the RG almost perpendicular to flow (Fig. 1) and are also just above the point where the glacier bifurcates. The 0248 elevation difference data indicate that RG may be losing more ice through a short, steep southern outlet into a bay within the Larsen B embayment west of Cape Fairweather. MODIS data indicate this area lost its shelf ice remnant (~44 km² in 2002, Fig. 1) during early 2004 and the ETM + imagery from 2008 show that the embayment extended well inland past the R&R-GL by that time. This loss of buttressing appears to have increased ice flow in the southern outlet, and may have caused an increase in the elevation loss rate in the ICESat profile in that sector relative to the main RG channel flowing toward the SNIS. Comparing the left side of Figure 5a to the right side shows that the southern side of the RG is losing elevation at 0.58 m a⁻¹ whereas the lower, northern side is losing elevation at 0.29 m a⁻¹. Further, the southern outlet's loss rate at Track 0248 is slightly greater than the mean loss rate for the main channel of the RG from the Track 0256 data (0.53 m a⁻¹, Table 4). From the available data, we cannot define when increased ice flow to the south began but the ETM + imagery series suggests that crevasses were more apparent in the ice moving south from the RG by 2008.

These ICESat data were not sufficient to calculate a basin-wide dH/dt. Consequently, the Scambos and others (Reference Scambos2014) analysis relied on DEM data to estimate net losses of 1.52 m a⁻¹ for the whole RG basin (see their Table S2, Basin 26a). The upper RG results from Track 0248 (both north and south) and the lower RG results from Track 0256 are both smaller than this area-averaged value, yet our results are not inconsistent with the overall basin analysis (Table 4). This is due to large loss rates derived from the DEM differencing results that extended the Scambos and others (Reference Scambos2014) analysis to areas not profiled by ICESat. Close examination of Figure 2 from Scambos and others (Reference Scambos2014) shows high losses (~8 m a⁻¹) for the lower part of the southern outlet of RG. A similar loss rate was derived for the lower part of a small unnamed glacier on the northern side of the overall RG basin area similarly exposed after the iceberg losses in 2004/05 from the north side of the SNIS. Though not definitive, the differential losses across Track 0248, with large losses from the published DEM-differencing plus more visible crevassing on the southern outlet seen in the imagery sequence, suggest this branch of RG may be decreasing the amount of ice reaching the SNIS.

Partitioning surface and basal melting

The combined RG and RI grounded ice elevation loss data suggest an approximate regional magnitude of the atmospherically-driven ablation rates for the SNIS region. With losses of more than 0.5 m a⁻¹ in the lower RG (Fig. 5b) and measured losses averaging 1.7 m a⁻¹ for the lowest ~70 m of the RI Track 0003 data (Fig. 4b, combining both north and south side RI data), we infer a magnitude of the atmospherically-driven thinning across the study area at elevations close to sea level as discussed below. Velocity data from Rack and others (Reference Rack, Rott, Siegel and Skvarca1999) is compatible with a more extensive compilation by Rignot and others (Reference Rignot, Mouginot and Scheuchl2011) and shows shelf ice velocities on the order of 25 m a⁻¹ or lower for the shelf areas suggesting that dynamic changes to the floating ice shelf are negligible. Table 4 shows the estimated thinning rates for the four ICESat tracks across the SNIS when the surface elevation change of this dense floating ice is converted to ice thickness. These data suggest that the floating ice thinning rate has a slight increase from west to east from the SNIS profiles.

We contrast the thinning estimated at Track 1349 (2.74 m a⁻¹) with the below 70 m elevation losses from nearby Track 0003 on RI (1.70 m a⁻¹) that has similarly low flow velocities (Rignot and others, Reference Rignot, Mouginot and Scheuchl2011). A ratio of the ice cap losses to the nearest ice shelf losses suggests that the atmospheric contribution to the shelf losses is ~62% of the total with ~38% due to subsurface oceanic melting. As a further test, we contrasted the estimated thinning in the lower RG below Track 0256 (Fig. 5b, Table 4 assuming it increases to the mean of the loss rates on the lower RG data and those below 70 m elevation on RI; the mean of 0.53 and 1.70 is 1.12 m a⁻¹) with the thinning at the Track 0367 location (2.39 m a⁻¹), which suggests that the atmosphere contributes 47% of the overall thinning and the ocean the other 53%.

In both cases, the estimated ocean contribution to shelf thinning is ~1 m a⁻¹, which is compatible with basal thinning reported further south in Holland and others (Reference Holland2015). It is important to note that the number of nunataks as well as a limited amount of adjacent bathymetry data (Lavoie and others, Reference Lavoie2015) both indicate that the water depths in the area are shallow (~300 m or less). This suggests that the basal thinning seen for the remnant shelf profiles may be due to tidal fluxing of a near-surface ocean layer, and not to incursion of deeper oceanographic water types. Similar behavior, i.e. increased basal melt inferred to be a result of tidal fluxing of solar-warmed surface water, has been reported for the McMurdo Ice Shelf ice front (Stern and others, Reference Stern, Dinniman, Zagorodnov, Tyler and Holland2013).