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New advances in geochronology and numerical modeling are revolutionizing our understanding of the interactions between climate, tectonics, and erosion. Nowhere are these interactions stronger than in alpine glacial environments, where erosion rates are higher and more spatially concentrated than in other mountain belt geomorphic process zones. At large scales, the efficacy of glacial erosion near the equilibrium line (i.e., the location where ice accumulation transitions from positive to negative) is thought to result in a climatically controlled limitation on elevation (Brozovic et al., 1997; Meigs and Sauber, 2000). At smaller scales, glaciers may increase relief through a combination of focused erosion in valleys and regional isostatic rebound (Montgomery, 1994; Small and Anderson, 1998). The expansion of glaciers beginning in the late Miocene has likely played a major role in the order-of-magnitude increase in erosion rates (inferred from sedimentation rates in adjacent basins) measured between 4 and 2 Ma in many regions worldwide (Peizhen et al., 2001). This hypothesis is supported by a number of recent low-temperature themochronologic studies that have dated the onset of accelerated late Cenozoic alpine glacial erosion from between late Miocene and Pleistocene time (e.g. Farley et al., 2001; Shuster et al., 2005; Ehlers et al., 2006). In this issue of Geology, two studies add new insights to the role of glacial erosion in mountain building. Berger and Spotila (p. 523 in this issue) use low-temperature thermochronology to provide the first direct test of whether glacier sliding rates exert the predominant control on glacial erosion rates over geologic time scales. Mevedev et al. (p. 539 in this issue) use numerical modeling to document the importance of iso-static rebound in producing rock uplift and relief in the Fjord Mountains of East Greenland, a Cenozoic passive margin.

In a pioneering study, Small and Anderson (1998) estimated the relief produced by glacial erosion (which is concentrated in valleys) and the subsequent isostatic response (which uplifts peaks and valleys, alike, within a region ~100 km wide, causing some peaks to increase in elevation). Small and Anderson used gently dipping and slowly eroding “summit surfaces” to reconstruct the paleotopography prior to the onset of late Cenozoic glaciation. The “missing mass” between the paleotopography and modern topography was then input into a two-dimensional flexural-isostatic model to infer how much of the modern relief could be attributed to glacial erosion. Because 80% of eroded bedrock is replaced isostatically, isostastic rebound can uplift plateau surfaces surrounding deeply incised canyons to elevations as much as four times their pre-isostatic levels (Pelletier, 2007). Similar studies in Antarctica (Stern et al., 2005) and the European Alps (Champagnac et al., 2007) have documented the importance of this isostatic relief-production process in glaciated landscapes. In this issue, Medvedev et al. tested the isostatic relief-production hypothesis in the fiords of eastern Greenland using methods broadly similar to the original work of Small and Anderson. Medvedev et al., however, extended the approach to three dimensions (3-D) and included the effects of sediment loading adjacent to the range. The 3-D pattern of erosional unloading can be very important in controlling the magnitude and spatial distribution of uplift (e.g., Pelletier, 2004), and sediment loading enhances rebound of the eroded range because the range becomes a peripheral bulge associated with loading-induced subsidence offshore. Although isostatic rebound can be responsible for significant relief production and rock uplift, the mean elevation of any range can only decrease during isostatic rebound. Therefore, although Medvedev et al. conclude that 1.1 km of late Cenozoic rock uplift occurred in this region, most of the 1.2 km of mean surface uplift that occurred following the deposition of Mesozoic marine sediments must have occurred prior to the onset of isostatically driven uplift. Medvedev et al. do not provide a mechanism for this surface uplift, but the thin crust of eastern Greenland suggests crustal thinning as one possible uplift mechanism, as has been suggested for western Greenland (Japsen et al., 2006).

Few direct measurements exist of the spatial distribution of glacial erosion over geologic time scales. For this reason, the results of Berger and Spotila (this issue) are exciting. Glacial erosion is mechanically complex, but theoretical models of glacial quarrying indicate that erosion is controlled by basal sliding speed (and is, perhaps surprisingly, independent of ice thickness) (Hallet, 1996). Hallet's sliding-speed–controlled glacial erosion model is commonly used in glacial landscape evolution models (e.g., MacGregor et al., 2000), hence a direct test of this model has fundamental implications for our understanding of glacial landscapes. Using the low-temperature thermochronometer (U-Th)/He in apatite, Berger and Spotila found the highest rates of exhumation on the windward (southern) side of St. Elias Range of southern Alaska to be located ~35 km inland from the coast, along a coast-parallel line that coincides with the mean Quaternary equilibrium line. Because sliding speed is greatest beneath the equilibrium line, Berger and Spotila's work corroborates Hallet's mechanical model of glacial erosion, and illustrates a direct climatic control on the spatial pattern of crustal deformation and rock uplift. Berger and Spotila suggest, however, that a model based on sliding speed alone is insufficient to explain their data because large and small glaciers were found to be equally erosive. Berger and Spotila argue that because ice discharge increases proportionately with drainage area, an erosion law based on sliding speed alone would predict that larger glaciers erode more rapidly. In fact, glaciers in larger catchments do not necessarily have greater sliding speeds. Although ice discharge is greater in larger catchments, larger catchments have wider and deeper glaciers that accommodate larger ice discharges even at similar sliding speeds. Available data on the scaling of glacier width and depth to area use glacier surface area, not catchment area, as the independent variable (Bahr et al., 1997). Nevertheless, when correlations between glacier width, depth, and surface area are taken into account, sliding speed can be expected to depend only weakly on glacier surface area (based on a global average). As such, the concentration of glacial erosion beneath the equilibrium line and the independence of erosion rate on catchment area documented by Berger and Spotila are both broadly consistent with a sliding-speed–controlled model for glacier erosion over geologic time scales.