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Glacial-geologic evidence from the Laurentide and Fennoscandian ice sheets indicates the presence of frozen beds, wet beds, and surficial meltwater at various times (e.g., Sugden and John, 1976); however, most of the data reflect processes near the ice margin during retreat. A likely history is that large areas of the Wisconsinan midlatitude ice sheets were frozen to their beds during growth. The bed probably thawed in central regions during the glacial maximum. Thawed beds existed beneath channelized ice flows across large areas of the southern margins and smaller areas of the northern margins, but frozen beds occurred between these channelized flows. Surficial meltwater reached the bed in large quantities in marginal areas, especially on the southern margin during ice-sheet retreat. (For a similar but more exhaustive review of much of this material, the interested reader is referred to Hooke, 1977; Sugden, 1977; and Hughes, 1981).
Carbon Dioxide in Glacial Meltwater The amount of CO2 available in ice for chemical weathering can be estimated reasonably well and is not large (e.g., Stauffer and Berner, 1978; Neftel et al., 1982; Oeschger and Stauffer, 1986). During the transformation of snow to ice without melting, the ice traps about 10 percent air by volume with the prevailing atmospheric CO2 concentration. This is retained in the air bubbles as CO2 (or in some form such that CO2 gas is recovered if the ice is crushed under vacuum), for a total of about 10-6 mole CO2/kg ice (CO2:H2O molar ratio ≈ 2 x 10-8, within a factor of two to three depending on atmospheric CO2 partial pressure, atmospheric pressure when air was trapped, and other variables). By comparison, rainwater in equilibrium with the atmosphere has a CO2 concentration an order of magnitude or more higher than this. Refrozen surficial meltwater in glaciers falls between the ordinary ice and rainwater CO2 concentrations.
When ice is melted beneath a glacier, water production rates are low, flow velocities are slow, and flow paths to the ice front typically are long; thus, it is likely that chemical weathering continues to completion and consumes the CO2 in the meltwater. Wet-based glaciers are reasonably efficient at eroding unconsolidated sediment in some areas and depositing it elsewhere (see below), so it is unlikely that organic-rich soils will remain in contact with a glacier and supply abundant CO2 to meltwaters over any significant time.
Under certain circumstances, surficial meltwater can remain in contact with both rock debris and atmospheric CO2. For example, compressive stresses near ice fronts may cause folding or thrusting of basal debris to the ablating ice surface. Under such open-system conditions, meltwater can obtain CO2 from the air as it is consumed in weathering; total CO2 consumption then can exceed initial CO2 content of the water (Raiswell, 1984). This process continues beyond the glacial margin in proglacial streams and differs from nonglacial streams only in that (1) glacial streams often transport material with more total area of fresh, unweathered mineral surface than similar nonglacial streams (see below); but (2) glacial streams generally lack CO2 input from organic-rich soils and thus have much lower CO2 concentrations than equivalent nonglacial streams.
Observations bearing directly on glacial chemical weathering are relatively scarce. Edmond (1973) and Hurd (1977) showed that Antarctic glaciers contribute very little to the dissolved silica in ocean waters adjacent to the Antarctic. Mountain glaciers are observed to have two meltwater systems: one with slow drainage and limited CO2 in which chemical weathering proceeds to equilibrium and solute concentrations are relatively high, and one with rapid, channelized drainage of surficial water in which solute concentrations are low, whether that drainage occurs on, in, or under the glacier (Collins, 1979; Raiswell, 1984). Proglacial mixing of these waters allows the CO2- bearing surficial meltwaters to weather silt supplied by the basal waters, but the extent to which equilibrium is reached and further CO2 is taken from the air is not known well. Ford (1971) found that glacial meltwater above tree line in limestone regions of the Canadian Rockies was saturated with respect to CaCO3 at only 50 to 90 mg/liter, compared to groundwater saturation at 100 to 265 mg/liter below tree line and creek and lake saturation at 100 to 140 mg/liter below tree line.
Glaciated basins in high mountains in maritime climates have high denudation rates (= 1 mm/yr) and chemical weathering is faster than the global average, but runoff is very high (4 m/yr is typical; Corbel, 1959; Reynolds and Johnson, 1972). No data are available on how much of that chemical weathering is associated with soil processes in unglaciated parts of the basins; the work of Ford (1971) suggests that much of it may be. In addition, glaciers in such maritime climates often have abundant englacial and supraglacial debris from rockfalls along with abundant meltwater; thus they are poor analogues for all except the narrow margin of a continental ice sheet.
Based on this discussion, chemical weathering in contact with continental ice sheets can be summarized as follows:
large areas of ice sheets are frozen to their beds and slow or stop chemical weathering, especially during ice-sheet growth;
central ice-sheet regions with thawed beds produce <O(1 mm/yr) basal meltwater, with CO2 available for weathering equivalent to that in <O(0.1 mm/yr) rainwater, and this CO2 generally is consumed in rock weathering;
fast-moving marginal regions produce O(10 to 100 mm/yr) basal meltwater over perhaps 10 percent of the ice