Short Name:
steffen_0732946

IPY: Stability of Larsen C Ice Shelf in a Warming Climate

Our ground penetrating radar surveys focused on key structural features of the ice shelf, including basal crevasses, marine ice and rifts. Here we present our published work on basal crevasses. Analysis of the marine ice and rifts is ongoing. We first studied a series of crevasses extending downstream from Churchill Peninsula. The key results were published in McGrath et al. (2012a). Twenty-seven basal crevasses, with an average spacing of 1.2 km, are identified along a 31 km transect aligned with the flow direction of the ice shelf downstream of Cabinet Inlet. Basal crevasses are identified by strong hyperbolic reflections within the ice column (Figure 1; black arrows) and two, nearly symmetric hyperbolas at the bottom corners of the crevasse (Figure 1; yellow arrows). The underside of the ice shelf is highly fractured in the initial 2 km of the transect, however, a smooth basal surface exists along the remainder of the transect (Figure 1). Crevasse spacing ranges from 0.5 to 2.0 km and does not systematically vary along the length of the transect. Penetration height of the basal crevasses is greatest near the origin, where they extend 120-134 m into the ice column (Figure 2). Height subsequently decreases along the length of the profile, decreasing by ~40 m in the first 15 km while remaining nearly constant over the remaining ~15 km (Figure 2). In addition to a decrease in absolute height, the relative penetration height (as a function of ice thickness) also decreases from ~40% to 25%. Initially, the bottom width of the basal crevasses is quite narrow (20-70 m) and at times, difficult to determine due to the overlap of the two hyperbolas that originate from the bottom corners of the crevasse (Figure 1A, yellow arrow). The opening width of the basal crevasses increases along the length of the transect, opening ~100 m in the first 15 km, while ranging between 150-240 m over the remaining ~15 km (Figure 2; black line). The surface undulations have an amplitude of 3-4 m, although the amplitude of these features are likely much larger proximal to Churchill Peninsula, as the radar transect was conducted near the edge of the features. Internal layering within the firn shows significant downwarping of 11-18 m above the basal crevasses, with deeper layers showing greater downwarping (Figure 1, red arrows). Surface undulations on the ice shelf are easily identified in both MODerate-Resolution Imaging Spectroradiometer (MODIS; 250 m) and Landsat (15 m) imagery as a series of parallel alternating dark and light bands (Figure 1). In total, 102 undulations, with the troughs aligned with the basal crevasses, are observed along a 155 km flow line transect. High-resolution commercial visible imagery (GeoEye; 1.65 m resolution) clearly shows a lack of surface crevassing upstream of the origin of the basal crevasses (Figure 3a). Surface crevassing, first observed as narrow (1-5 m) bridged features, is apparent where the first surface undulations are observed (Figure 3b). The surface crevasses become increasingly abundant and well defined in the along flow direction, reaching widths of 8-24 m, while remaining primarily aligned with the crests of the surface undulations (Figure 3c). In our final field season, we did a focused radar survey of one isolated basal crevasse and the subsequent surface expression of this feature. The key results from this work are summarized in McGrath et al. (2012b). Visible imagery details a series of isolated surface depressions in the main outflow of Cabinet Inlet, extending seaward to the calving front (Figure 4a). We focus our observations on one of these features (Figure 4b). The surface depression has a maximum depth of 13.0 m and extends for 4.5 km, as measured from kinematic GPS and imagery, respectively (Figure 4b). A large hyperbolic reflection within the ice column is aligned with the surface depression, which we interpret as the apex of a basal crevasse (Figure 2a). It extends 233 ± 11 m in height from the base of the ice shelf, penetrating through over 66% of the mean local ice thickness and is 470 m in width (Figure 5a and b). The size of the basal crevasse increases the local ice-ocean interface by ~30% relative to a flat-bottomed ice shelf. Above the hyperbolic reflection, the firn and upper ice layers down warp by 15-20 m (Figure 5a). Numerous snow-bridged surface crevasses, with widths between 20-25 m, are aligned parallel to the basal crevasse but on the down sloping flanks of the surface depression, as observed in both the visible imagery (Figure 4b; shadowed features indicated by red arrow) and in the hyperbolic reflections in the upper 10 m of the radargram (Figure 5a; highlighted in red). We attribute the formation of the surface crevasses to bending stresses induced as the ice shelf surface deforms in order to reach a modified hydrostatic equilibrium [McGrath et al., 2012]. We offer the following observations here to support this hypothesis. We observe two nascent basal crevasses, which are located upstream from the series that extend seaward from Churchill Peninsula, which have propagated to a similar height but have limited surface deformation (1-4 m) and no apparent surface crevassing (Figure 6). Further along flow (i.e. with increased temporal evolution), the basal crevasses have clearly defined surface depressions, down warped firn and ice layers above the basal crevasse and numerous surface crevasses adjacent to the basal crevasse [McGrath et al., 2012]. Together, these observations suggest that the basal crevasse forms first, subsequently followed by surface deformation, the visco-elastic response to the reduced ice thickness above the basal crevasse. However, as this thinner section is partially supported by the full thickness ice to either side, the resulting geometry induces bending stresses, with tension across the crests and down the flanks, sufficient to induce surface crevassing, and compression in the surface depressions, which are free of surface crevasses (Figure 1b). The significance of basal crevasses, especially in the context of meltwater ponding and hydrofracture is highlighted below. Meltwater-driven hydrofracture, the process by which water filled surface crevasses fracture downwards, has been suggested to be an important mechanism in the final break up of several ice shelves [Weertman, 1973; van der Veen, 1998, 2007; Scambos et al., 2000; 2003]. We have shown that basal crevasses, beyond introducing large-scale ice shelf weaknesses, can create both surface depressions and surface crevasses. The most apparent implication of meltwater ponding in the surface depression is if the meltwater were to intersect a flanking surface crevasse, and subsequently establish a channel by which the pond could drain, thereby providing the necessary water volume for continued fracture. Perhaps less obvious, however, is that the increased load in the trough will increase extensional stresses along the flanks and in the vicinity of the basal crevasse apex, potentially leading to further propagation and the possibility for a shear fracture to connect these features [Bassis and Walker, 2012]. This structural weakness could be further exploited if hydrofracture originates from the base of the surface trough, where the hydrostatic pressure is the greatest, and where, despite the large-scale compressional environment, incipient surface cracks / fractures are still likely to be present [Fountain et al., 2005]. The presence of the basal crevasse greatly reduces the ice thickness in the vicinity, thereby minimizing the distance through which these small fractures have to propagate prior to creating a full-thickness rift. While the surface crevasses certainly weaken the ice shelf, this latter case highlights the possibility that it is the presence of the basal crevasse that is more important for ice shelf stability. Basal crevasses are an order of magnitude larger in width and depth than the surface crevasses they create, and by concentrating meltwater ponding directly above them, they can control fracture location, and therefore, ice shelf disintegration. In addition to the observations of melt pond drainage on Larsen B, sediment cores retrieved from beneath both the former Larsen A and Prince Gustav Ice Shelves, record spatially discrete sediment pulses interpreted as the drainage of supraglacial lakes and/or crevasses prior to the ice shelf disintegration event [Gilbert and Domack, 2003]. Together, these observations provide clear evidence that fractures do propagate through the ice shelves, although the location where the hydrofracture originated is unclear (i.e. whether it was a proximal surface crevasse or incipient flaw beneath the pond). A corollary can be drawn to supraglacial lake drainage on the Greenland Ice Sheet, where fractures, and later moulins, develop within the lake boundary [Das et al., 2008]. Thus, if hydrofracture does originate from within the pond boundary, the presence of the basal crevasse should make the formation of a full-thickness rift exceedingly efficient. Analysis of the automatic weather station data is ongoing and is focused on conditions that are conducive to surface meltwater production. Particular attention is being paid to föhn events, in which air that passes over the peninsula is adiabatically warmed, and leads to meltwater production on Larsen C. The climate of the peninsula has been relatively cool over the 4 years of the project. We considered this recent variability in the context of the 30-50 yr climate trends in McGrath and Steffen, 2012. Atmospheric warming over the Antarctic Peninsula (AP) during the second half of the twentieth century was remarkable, with local trends exceeding 0.5°C decade-1 (Vaughan et al. 2003; Turner et al. 2005). The complexity of atmospheric and oceanic circulation around the AP precludes clear attribution of this warming, although modified circulation patterns, largely driven by changes in latitudinal pressure gradients measured by the Southern Annular Mode (SAM) index, likely played an important role (Vaughan et al. 2001; Thompson and Soloman, 2002; Marshall 2003, 2007; van den Broeke and van Lipzig, 2004). There has been increased advection of warm, maritime air towards the AP, coupled with increased upwelling of Circumpolar Deep Water onto the continental shelf of the Amundsen-Bellingshausen Sea (ABS; Orr et al. 2008; Jacobs et al. 2011; Martinson 2011). Dramatic reductions in sea ice extent (-6.6% decade-1) in the ABS likely contributed to this warming as well, possibly as a driver and certainly as a positive feedback (Meredith and King 2005; Turner et al. 2009; Stammerjohn et al. 2012). In contrast to the strong warming trends observed during the previous half-century, between 2000 and 2011, AP stations have shown much greater variability, with many stations showing a slight cooling trend (Fig. 7). The mean annual temperature trend at Marambio station switched from a warming of 0.65 ± 0.47°C decade-1 between 1971 and 2000 to a cooling of -1.57 ± 2.4° C decade-1 between 2000 and 2011. The cooling has been strongest at other northern AP stations (Bellingshausen, O’Higgins, Esperanza), with a cooling trend of about -0.7°C decade-1 over the last 15 years, following a tendency that started with the end of the 1998 El Niño. However, marked cooling is also evident on the Larsen C Ice Shelf, the largest remaining ice shelf on the AP, located on its eastern edge. This site has cooled by -1.11 ± 2.4°C decade-1 between 2000 and 2011, although interannual variability is large. It is important to note that none of the reported trends over the past decade are statistically significant, and cooling of a similar magnitude previously occurred over short time periods earlier in the record (Fig. 7). However, with the observed variability over the past decade, there is no longer a statistically significant (p

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