During the expansion of Tasman Lake and retreat of Tasman Glacier between 1990 and 2006, downwasting of the lower terminus continued and resulted in the extensive formation of supraglacial ponds evident in Figure 4.1A–C. Pond development and growth appears to have been centralised in the area of greatest depth as evident in the 2008 bathymetry of Tasman Lake (Figure 4.9). As surface topography of a glacier typically mirrors bed topography (Benn and Evans 2010), the development of supraglacial ponds in this area is likely to be related to the surface expression of this depression in the glacier bed providing a loci for the pooling of meltwater and the collapse of englacial conduits (Benn et al., 2001; Röhl, 2006, 2008). This coupled with the decrease in ice velocities between 1995 and 2002 (Röhl, 2005) and the low gradient (< 1.5°) nature of the lower glacier tongue between 1990 and 2006, provided all the necessary conditions for pond development (Sakai et al., 2000; Benn et al., 2001; Röhl, 2008).
Once hydraulically–connected to the main drainage network (either englacially or to the main lake), the rate of pond growth accelerated between 2003 and 2006 as a result of enhanced subaerial and subaqueous melt and calving (Sakai et al., 2000; Benn et al., 2001; Röhl, 2008). Decreased ice velocity of the lower glacier between 2004 and 2006, evident in Redpath (2011, p. 115), also appears to correspond to the expansion of supraglacial ponds over this period of time. By 2006 supraglacial ponds on the lower glacier had expanded to 0.46 km2, including ponds partially connected in planform to Tasman Lake. Once connected to the lake, increased water temperature and circulation within the ponds dramatically increased the rate of ice loss from the terminus of Tasman Glacier, significant altering terminus dynamics (Röhl, 2008).
Figure 4.9: Bathymetric map of Tasman Lake with the position of the Tasman Glacier terminus and the extent of supraglacial ponds in April 2000 and January 2006. In general, pond growth is in the region of greatest water depth of Tasman Lake in 2008.
It was this rapid expansion of supraglacial ponds and retreat of the glacier down a reverse slope into deeper waters (Figure 4.7) that led to the disintegration of a large section of the lower terminus between 2006 and 2007 and the loss of 1.47 km2 of ice (Table 4.1). Figure 4.10 and Figure 4.11 show the terminus of Tasman Glacier before and after this event. They indicate that a large planform area of formerly debris–covered ice in the January 2006 satellite image (Figure 4.1D) had been lost by pond and lake expansion and calving by March 2007 (Figure 4.10A). As pond growth primarily occurs in a horizontal direction (Röhl, 2008), the planform area lost is assumed to have remained ice cored as large sections of the glacier visible in Figure 4.10A remained in situ up until the entire section disintegrated (Figure 4.11). The loss of ice via the expansion of ponds on the lower glacier would have significant increased buoyant forces acting at the terminus zone as the substitution of water for ice within depressions decreases the ice overburden pressure in the order of 10 tonnes m-2 (Röhl, 2008). Pond expansion also had the effect of decreasing ice thickness of a large area of the terminus to below lake level. Retreat of the glacier into progressively deeper waters occurred between 2000 and 2006 (Figure 4.7), with
the greatest water depths recorded in the bathymetric survey in 2008. Hence, in the area of ice loss in 2007 (Figure 4.1), increased torque arising from buoyant forces (Dykes et al., 2011) across this section of the terminus would have been significant. This is important because as grounded fresh–water glacier termini become buoyant and attain flotation, the rate of ice loss increases substantially as a result of changes in glacier dynamics and stress regimes (Warren et al., 2001; van der Veen, 2002; Röhl, 2008). It was the increase in stresses from applied buoyant forces that led to the disintegration of large section of the terminus during the start of April 2007 (Figure 4.11). This event produced numerous large icebergs (Figure 4.10B) that calved coherently from the terminus.
Figure 4.10: Oblique aerial photography of the terminus of Tasman Glacier in (A) March 2007 showing the state of the terminus immediately prior to breakup (Photo: S. Winkler) and (B) February 2008 after breakup. Note that several large and coherent icebergs have already calved from the terminus in (A) and that a large number of icebergs were preserved over the following year.
Previous work (e.g., Dykes et al., 2011) has indicated that terminus disintegration in April 2007 was potentially initiated by significant rainfall falling within the region in the months prior to calving. Rain falling within the Tasman Glacier catchment can cause up to 5 m increase in lake level over short (c. 48 hours) time periods (see chapter 5). Such fluctuations at an unstable terminus have the potential to instigate calving by exceeding the threshold of flotation (Boyce et al., 2007). The direct link between lake level fluctuations and buoyancy–driven calving at Tasman Glacier remains unclear as not all events of this type are associated with increased lake–level (chapter 8). However, at a terminus approaching the critical threshold for flotation, lake–level fluctuations may increase stress into the terminus region, further destabilising the terminus and increasing the likelihood of high–magnitude calving events. This is discussed further in the following sections.
Figure 4.11: False colour ASTER image of Tasman Lake at the start of April 2007 showing the disintegration of the lower section of the Tasman Glacier terminus pictured in Figure 4.10A. The presence of white icebergs within the region of the former terminus indicates that calving has taken place in the days prior to image acquisition.