5.1. Pathways for Organic Carbon Turnover
 A number of functional microbial types have been identified in subglacial microbial consortia, including aerobic heterotrophs [Skidmore et al., 2000], nitrate reducers [Skidmore et al., 2000; Kivimaki, 2004], sulphate reducers [Skidmore et al., 2000; Kivimaki, 2004], Fe(III) reducers [Tung et al., 2006] and methanogens [Skidmore et al., 2000; Tung et al., 2006]. The production of methane from these consortia [Skidmore et al., 2000; Tung et al., 2006] also implies that syntrophic bacteria are present, since these function in concert with methanogens to produce methane. Methanogenesis represents the last step for organic matter degradation, and takes place in the most anoxic environments once oxygen, nitrate, sulphate and Fe(III) have been removed by reduction.
 Redox reactions in glacial environments can take a number of different pathways, with the reaction products of one set of reactions often being used in other reactions, so driving Eh down. For example, oxygen is often used to oxidize sulphides in glacial environments at relatively high Eh, and the reaction product, sulphate, can be used to oxidize organic matter in the absence of oxygen at lower Eh. Ultimately, organic matter is oxidized, to a degree that is controlled by the initial concentration of oxygen. We estimate how much organic carbon may be oxidized in subglacial environments during aerobic and anaerobic respiration by assuming that the different terminal electron acceptors directly oxidize organic matter (rather than react with bedrock components, so producing compounds which, in turn, may oxidize organic matter).
 First, we calculate the amount of organic carbon degradation (OCD) in the SOC that would occur (equation (7)) assuming (1) that the only oxygen supply to the bed is from geothermal melting of basal ice and (2) that the oxidation of organic carbon is the first sink for this oxygen (as opposed to other reactions that consume atmospheric oxygen, such as sulphide oxidation). The rate of organic carbon degradation and carbon dioxide production (in moles) is equal to the rate of oxygen liberated from ice melt (in moles):
The rate of oxygen production, OP (in mol m−2 ka−1), can be calculated as follows:
where VM is the volume of meltwater produced by geothermal heating of basal ice (in m3 m−2 ka−1), GC is the typical gas content of basal ice (0.1 mol m−2 ka−1; derived from gas concentrations of 0.1 cm3 g−1 [Souchez et al., 2003]), F is the fractional percentage of oxygen present in the atmosphere, and G is the universal gas constant (22.4 L mol−1).
 The amount of oxygen liberated by the melting of basal ice is dependant on the basal melt rate. Melt rates of 5–6 mm a−1 are typically used as representative values for the rate of basal melting beneath the Northern Hemisphere ice sheets [Breemer et al., 2002], although lower area averaged values have been computed (2–4 mm a−1 [Clarke et al., 2005]). These values compare well with maximum contemporary rates for the Greenland Ice Sheet (7 mm a−1 [North Greenland Ice Core Project Members, 2004]). We use the upper of these values (6 mm a−1) in our calculation. Melt rates of 6 mm a−1 give values of OCD values of 67.8 g C m−2 ka−1. Calculated over the areas of the Laurentide/Innuitian/Cordilleran and European ice sheets over 75 ka of glaciation (warm-based area of 30%) and 10 ka of deglaciation (warm-based area of 60%), we derive a value of 35 Pg C for potential OCD by aerobic microbial respiration during the last glacial cycle (Table 5).
Table 5. A Summary of the Amount of Organic Carbon Degradation That Takes Place Aerobically and Anaerobically at the Ice Sheet Base and, by Extension, the Amount of Carbon Dioxide Produceda
 The full utilization of oxygen during the respiration of organic carbon promotes anaerobic respiration, when compounds such as nitrate, sulphate and Fe(III) are reduced. These compounds can arise from ice melt and glacial flour. Ice melt is a source of sulphate and nitrate, with concentrations typically of the order of ∼1.0 μeq L−1 [Röthlisberger et al., 2002; Herron and Langway, 1985]. Nitrate may be reduced during denitrification or nitrate reduction (equations (7) and (8), respectively):
Sulphate is reduced according to equation (11), in which 2 mol of organic carbon are oxidized for each mole of sulphate reduced:
The amount of organic carbon degraded anaerobically (OCDAN) during sulphate/nitrate reduction is calculated as follows:
where VM is the volume of meltwater produced by geothermal heating of basal ice (in m3 m−2 ka−1 and assuming a basal melt rate of 6 mm a−1), R is the molar ratio of nitrate/sulphate reduced:organic carbon oxidized, C is the concentration of sulphate or nitrate (in mmol m−3), A is the combined area of the Laurentide/Cordilleran/Innuitian and European ice sheets and t is time (ka). We calculate that 0.11–0.18 Pg C and 0.18 Pg C would be oxidized to carbon dioxide during nitrate reduction and sulphate reduction, respectively, under ice sheets during 75 ka of glaciation and 10 ka of deglaciation.
 The primary source of Fe(III) is glacial flour. Fe(III) can be used to oxidize organic matter, as the following equation shows (equation (13) [after Stumm and Morgan, 1996])
The average ratio of Fe2O3:FeO in crustal rocks is ∼1:1 by weight, ranging from 1:1.1 for igneous rocks to 1: 0.7 for shales [Garrels and Mackenzie, 1971]. The ratio of Fe(III):Fe(II) in crustal rocks is ∼1:1.1, and so the ratio of Fe(III) to total Fe is ∼1:2.1. Let us assume that there is a 1m thick subglacial debris layer, and that this debris layer has a porosity of ∼30%. The density of siliclastic material is ∼2700 kg m−3, and hence the mass of siliclastic material in the layer is 1890 kg m−2. The mass of bioavailable Fe in glacial debris is 0.93% by weight (R. Raiswell, personal communication, 2007), and hence the mass of bioavailable Fe in the siliclastic debris is 17.6 kg m−2 and the mass of bioavailable Fe(III) is ∼8.4 kg m−2. According to equation (12), 4 mol (223.4 g) of Fe(III) oxidizes 1 mol (12 g) of organic C, and so 8.4 kg m−2 of Fe(III) will oxidize 0.45 kg m−2 of organic C during the last glacial cycle. Calculated over the areas of the Laurentide/Cordilleran/Innuitian and European ice sheets, this gives 6.5 Pg OC oxidized by Fe(III) during 75 ka of glaciation and 10 ka of deglaciation.
 A summary of the amount of organic carbon degraded aerobically and anaeobically, and by implication the amount of carbon dioxide produced, is presented in Table 6. The total amount of OC degraded by these means is equal to ∼42 Pg C, producing an equivalent amount of CO2 (in Pg C). These results produce higher estimates of aerobic organic decomposition than similar calculations performed by Skidmore et al. , who estimated that 8.1 Pg C would be converted to CO2 over a glacial cycle. If this latter estimate if the more correct of the two, the potential for conversion to methane only increases. We also produce much lower estimates of anaerobic degradation of organic carbon compared to Skidmore et al.  (up to ∼8000 Pg C over a glacial cycle), since we do not base our calculations on rates of sulphate reduction in marine sediments, where there is a plentiful supply of sulphate.
 The total amount of noneroded, bioavailable SOC that remains under the ice sheets for 38–75 ka and is available to methanogenesis is 125 Pg (167 Pg C – 42 Pg C degraded by aerobic/anaerobic respiration). A maximum of 50% of the 125 Pg C can be converted to methane according to the stoichiometry of methanogenesis by acetate reduction (equations (1)) and carbon dioxide reduction (equation (2)), in the latter case due to the limitation on hydrogen supply from fermentation (equation (14)):
These arguments suggest that 63 Pg C could potentially be converted to methane.
5.2. Rates of Carbon Turnover
 An implicit assumption of the reasoning in sections 5.1 and 5.2 is that the metabolic rates of subglacial microbes are consistent with them being able to turn over ∼63 Pg C to methane and 42 Pg C to carbon dioxide during a period of 38–75 ka. The following discussion explores this assumption, through a careful analysis of potential rates of subglacial microbial metabolism.
 The metabolic rate of microbes depends on a number of factors, to include carbon/nutrient supply, water and substrate availability and temperature. While at the pressure melting point and underlain by soil or till, the subglacial hydrology of the Laurentide and European ice sheets would comprise water flow through porous sediments [Walder and Fowler, 1994]. In sedimentary basins, recharge by glacial meltwaters would have also stimulated the development of deeper and continental-scale groundwater reservoirs [Boulton et al., 1995; Breemer et al., 2002; Person et al., 2007]. Thus microbes would be free to move and grow and there would be exchange of nutrients and organic carbon between different components of the subice sheet hydrological system. During the initial phases of glaciation, we believe that recent sequestration of organic carbon would result in a relatively well nourished microbial ecosystem with rates of metabolic activity depressed only by the low in situ temperatures. Rates of microbial metabolism would be expected to decrease over time as labile organic carbon compounds become depleted.
 No published rates of in situ microbial activity exist for subglacial environments. We draw upon rates derived for microbes in analogous environments, in the absence of direct measurements of subglacial microbial metabolic rates. None of these give a true indication of rates in the subglacial environment, since nutrient/carbon supply and moisture conditions are slightly different. Using this approach, however, it is possible to produce an envelope of metabolic rates within which we believe subglacial microbial activity would fall.
 During the initial phases of glaciation, and while relatively labile organic compounds are present, rates of microbial metabolism might be consistent with those observed in other nutrient/carbon rich “hydrologically open” environments. Examples of such environments might include the permafrost active layer, polar snow and firn. Carbon turnover rates of 10–102 g C g C−1 a−1 at 0°C [Price and Sowers, 2004] are found in such systems. There are no ideal analog environments for subglacial microbial activity once the organic carbon substrate has been significantly degraded. As a minimum estimate, we might consider deep ocean environments where limited carbon supply results in low rates of carbon turnover (10−6 g C g C−1 a−1 [Price and Sowers, 2004]). Although there are several rates now published for microbial metabolism in glacial ice, to include accretion ice of Lake Vostok, silty basal ice of the GRIP and GISP2 ice cores and other polar ice cores [Price and Sowers, 2004], we do not consider these as representative of the wet-based parts of the Laurentide and European ice sheets. Microbes inhabiting liquid water veins in ice have very low inferred rates of microbial activity (10−6 to 10−9 g C g C−1 a−1 [Price and Sowers, 2004]), indicating the metabolism of immobile, probably dormant communities or at best, metabolism of communities with low nutrient/carbon levels. These low rates of metabolic activity are in part a reflection of the fact that microbes have no or low mobility, which limits growth and metabolism [Price and Sowers, 2004; Tung et al., 2006]. This is not a constraint for microbes inhabiting a till layer beneath an ice sheet, which would have water saturated debris, an initially high carbon/nutrient supply and opportunities for renewal of small amounts of nutrients and carbon from basal ice melt thereafter.
 We produce first-order upper and lower estimates of potential methane and carbon dioxide production from subglacial environments beneath the North American and European ice sheets over the last glacial/interglacial cycle (75 ka of glacial conditions and 10 ka of interglacial conditions) using some of the metabolic rates presented above for analog environments. Calculations are given in Appendix A and results of the total mass of C produced as CH4 in Table 6. Assumptions made in these calculations are also presented in Table 6. These calculations are performed as follows,
 1. For the upper estimate, we assume that subglacial microbes (concentrations of 107 mL−1 [Sharp et al., 1999; Price and Sowers, 2004]) are able to metabolize carbon at rates similar to those displayed by permafrost microbes at 0°C when provided with a suitable carbon substrate.
 2. For the lower estimate, we assume that subglacial microbes have rates of metabolism in line with other parts of the deep cold biosphere (e.g., deep ocean sediments, Lake Vostok accretion ice and polar ice cores; 10−6 g C g C−1a−1) at 0°C and are present in concentrations of 107 cells mL−1. We emphasize that this estimate is unrealistic since advancing ice sheets will have much higher concentrations of labile subglacial organic carbon than those found in deep ocean environments and do not have the physical constraints on growth observed in liquid water veins in ice. We include this rate, however, for comparative purposes, and in acknowledgment that there are currently no measurements of the rate of subglacial carbon turnover under in situ conditions.
 These calculations give total estimates of carbon turnover (and gas production as CO2 and CH4). We subsequently partition the total carbon turnover estimates into the CO2 and CH4 components using three different methods. These methods are as follows:
 In method 1 we subtract our estimate of total CO2 production by aerobic/anaerobic respiration (42 Pg C) from the total carbon turnover (in Pg C). The remaining quantity is reduced by 50% in order to account for the incomplete conversion of SOC to methane either by acetate or CO2 reduction (equations (1), (2), and (14)). We assume that any OC not converted to CO2 is converted to CH4.
 In method 2 we employ a mass fraction of CH4:(CH4 + CO2), equivalent to that produced in incubation experiments by Skidmore et al.  (ratio of 0.54), to partition our total carbon turnover estimates.
 In method 3 we use a mass fraction for CH4:(CH4 + CO2), equal to that observed in the basal sections of the GISP2 ice core [Souchez et al., 2006] (ratio of 0.1), to separate the total carbon turnover estimates.
 We consider methods 1 and 2 as the most likely scenarios, since method 3 is based upon the ratios of carbon dioxide and methane produced in ice veins, where carbon supply and physical space for microbial growth/movement are not an appropriate analog for a water saturated till layer beneath an ice sheet. Methods 1 and 2 give very similar values for methane production (2320 and 2500 Pg C, respectively).
 These calculations give a broad spectrum of organic carbon turnover estimates, ranging from a total of 22–4680 Pg C, a reflection in itself on the uncertainty in rates of subglacial microbial metabolism and the partitioning of total carbon fluxes as gas into methane and carbon dioxide. We believe that the most realistic carbon turnover value for a wet-based ice sheet would lie somewhere between these two estimates, being of the order of 102 Pg C. We might expect 10–102 Pg C of methane production during the last glacial cycle depending upon which partitioning method we use for carbon dioxide and methane. Our estimate of SOC available for conversion to methane over 38–75 ka (63 Pg C) falls toward the low end of this estimate, suggesting that such a conversion might not be limited by subglacial microbial metabolic rates.
 This discussion highlights very clearly that detailed studies of in situ rates of subglacial microbiological activity are required in order to evaluate whether subglacial production of methane could be significant. Until the rates and pathways of subglacial microbial carbon cycling are better constrained, it is difficult to make firm conclusions regarding the potential for subglacial methanogenesis.