Thermokarst basins are closed depressions formed by degradation of ice-rich permafrost. They are generally 0.5-20 m deep and 0.01-5 km in diameter, and many contain standing water (thermokarst ponds and lakes). The basins are initiated by factors such as water ponding or vegetation degradation. Thermokarst ponds or lakes sometimes develop at sites where thaw occurs beneath standing water, notably at ice-wedge intersections or in low-centred polygons, as well as under small streams (Dredge and Nixon 1979). The likelihood of such site-specific disturbances may be increased by regional disturbances such as climate warming (Burn and Smith 1990).
Basins grow by deepening and widening. Deepening is promoted by ponding of water on the basin floor, especially if water depth exceeds the maximum thickness of winter lake ice (~2 m); when this occurs, the bottom-water temperature exceeds 0°C all year, resulting in continuous thaw of underlying excess ice and subsidence of the lake floor. Ponds and lakes also thaw permafrost around their margins, causing bank subsidence, slumping of lake shorelines and submergence or tilting of vegetation (Burn 1992). In lakes with sufficient fetch, wave-induced currents and lake-ice scour erode shores, remove newly thawed sediment or initiate thaw slumping (Rampton 1974). In central Yakutia, large thermokarst basins with steep sides and a flat, grass-covered floor (alases) develop by slumping of thermokarst mounds on basin margins and then by thermokarst subsidence beneath a thermokarst lake (Czudek and Demek 1970; Soloviev 1973). Some alases are several thousand years old, whereas others have formed during the span of a human generation. Alases may eventually coalesce, forming thermokarst valleys.
Limited data are available on rates of lake-basin enlargement. Wallace (1948) estimated bank retreat of ~ 0.06-0.18 m per year at two thermokarst lakes in eastern Alaska, and Burn and Smith (1990) employed comparison of aerial photographs taken in 1949 and 1984 to derive a mean growth rate of 0.7 m per year for 12 thermokarst lakes in boreal forest near Mayo. Radial expansion rates of 1.5-5.0 m per year (but accelerating through time) were determined for a high-mountain thermokarst lake on the Gruben rock glacier in the Swiss Alps (Kaab and Haeberli 2001).
Basin size and shape are controlled largely by the distribution and volume of pre-existing excess ice, the time since thaw commenced, and by erosion and sedimentation. Shallow lake basins, usually no more than a few metres deep, form by thaw of the ice-rich layer in near-surface permafrost (Sellmann et al. 1975), whereas basins 10-40 m deep represent thaw of much thicker ice-rich permafrost (Czudek and Demek 1970; Carter 1988; Romanovskii et al. 2000).
Basin growth may cease by lake drainage, infilling with sediments and peat, or exhaustion of ground ice (Burn 1992). Lake drainage is sometimes rapid. On the Tuktoyaktuk Peninsula, on average two lakes drain catastrophically each year (Mackay 1988). Drainage results mainly from diversion of water through interconnecting ice-wedge systems, causing rapid thermal erosion. Lake drainage is often incomplete, leaving shallower lakes or residual ponds. Basins infill by lacustrine or colluvial sedimentation, hydroseral encroachment by plants such as sedges and Sphagnum, peat accumulation and eventually growth of ice wedges. The sediments within them represent the most widespread type of thermokarst sediments, with a high preservation potential and often a distinctive stratigraphy (Hopkins and Kidd 1988; Murton 1996).
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