The snowball hypothesis has properly been questioned and the sharpest questions to date concern the terrestrial glacial regime, the isotopic evolution of snowball seawater, and the fate of eukaryotic organisms. These and other questions highlight the central difficulty with the hypothesis, which is our limited conception of a snowball event itself. With postulated conditions lying far outside familiar parameter space, there is danger that the hypothesis we question is a caricature of a snowball Earth.

Would glaciers flow?
Geological evidence for the existence of dynamic glaciers is incontrovertable. Many LNGD contain faceted and striated clasts (Fig. 2b), some far-travelled, and deformation structures caused by glacial flow (Cahen, 1963; Eisbacher, 1981; Wang et al., 1981; Edwards, 1984; Deynoux, 1985; McMechan, 2000). Associated sub-glacial pavements (Fig. 2a) are indelibly shaped, scratched, cracked, and polished by glacial motion (Reusch, 1891; Perry and Roberts, 1968; Deynoux, 1980; Rice and Hofmann, 2000). Are these observations compatible with the cold, dry atmosphere and limited hydrologic cycle of a snowball Earth (Christie-Blick et al., 1999; McMechan, 2000)? A major difficulty with this question is that we do not know which phase of the glacial cycle (Fig. 7) the deposits represent. A strong bias toward the end-phase is expected because of the temperature trajectory and because there would have been ample time for ice to thicken even if net accumulation rates were very low. The magnetic reversal record in the Elatina (Sohl et al., 1999), however, suggests that a significant time interval is represented.

Climate models suggest that an ice-albedo runaway might occur very rapidly (Hyde et al., 2000), freezing the tropical oceans before ice sheets had time to develop on low-latitude continents (Hoffman et al., 1998b; Pollard and Kasting, 2001). During a snowball event, mean temperatures slowly rise (Fig. 7) in response to atmospheric CO₂ and this will cause an exponential increase in saturation vapour pressures. Moisture sources include subaerial volcanic outgassing, sublimation of tropical sea ice (Warren et al., 2002), and evaporation of transient melt ponds (Walker, 2001), leads and polynyas within the tropical ice pack (Kirschvink, 1992). Given large seasonal and diurnal fluctuations in surface temperature (Walker, 2001), water vapor would be advected to topographic highs on summer afternoons and moisture would condense there forming glaciers (Walker, 2001). Due to the weak hydrologic cycle, snow accumulation rates would be extremely low, but a positive feedback would develop between surface elevation and snow accumulation if glaciers thickened. Over sufficient time, these glaciers must become wet-based and flow, possibly in localized ice streams or as episodic ice surges (Bentley, 1987; MacAyeal, 1993; Condon et al., 2002). The observed zonal distribution of LNGD (Fig. 1) might broadly reflect the zonal pattern of precipitation minus ablation on the snowball Earth. On the regional scale, the importance of topography, particularly coastal topography, in capturing the limited precipitation is not inconsistent with the common observation that thicker glacial deposits occur in tectonically more active areas and times (Young, 1995). They commonly overlie pre-glacial strata with angular unconformity, consistent with a prolonged hiatus (with ongoing block rotation) between the ice-albedo runaway and the onset of glacigenic sedimentation (Fig. 7). However, LNGD are missing or very thin beneath cap carbonates in many areas (Wright et al., 1978; Preiss, 1985; Aitken, 1991; Hoffmann and Prave, 1996), or aeolian sandstone overlies permafrost regolith in place of LNGD (Deynoux et al., 1989; Moussine-Pouchkine and Bertrand-Sarfati, 1997; Williams, 1998). As lag deposits or other evidence of erosion of LNGD before cap-carbonate deposition is generally lacking, the observed stratigraphic relations indicate that ablation exceeded precipitation through the glacial cycle in many areas. Where LNGD are absent, without evidence for removal by erosion, or where their position is occupied by aeolianite (Deynoux et al., 1989; Moussine-Pouchkine and Bertrand-Sarfati, 1997; Williams, 1998; Fig. 5a), ablation evidently exceeded precipitation over the full glacial cycle.

Would seawater change?
The elemental and isotopic composition of seawater should change in a snowball event. Decimated organic productivity and hydrothermal dominance should cause δ¹³C and ⁸⁷Sr/⁸⁶Sr, respectively, to fall (Kump, 1991; Jacobsen and Kaufman, 1999). The trajectories of the changes depend on the C and Sr residence times, and are sensitive to buffering by carbonate dissolution (Jacobsen and Kaufman, 1999) in response to submarine volcanic exhalations of CO2 and other acids. Consequently, few δ¹³C or ⁸⁷Sr/⁸⁶Sr values have been reported for the glacial ocean because primary syn-glacial carbonates are rare. Kennedy et al. (2001b) report positive δ¹³C values from purported examples in Australia, Namibia and California, whereas Xiao et al. (2001) find that thin calcareous crusts in the Tereeken glacials (Table 1) of NW China have δ¹³C close to -5‰.

Cap carbonates have been substituted as proxies for seawater ⁸⁷Sr/⁸⁶Sr during snowball events (Jacobsen and Kaufman, 1999; Kennedy et al., 2001b), justified by the Sr residence time of ~3.5 myr in the present ocean. Neoproterozoic Sr isotope data are not very reliable because strontium is a mobile trace element and most carbonates of that age are totally recrystallized (Derry et al., 1992). Nevertheless, the “least altered” ⁸⁷Sr/⁸⁶Sr ratios in cap limestones (dolomites have very low Sr concentrations) are not significantly different from pre-glacial values (Jacobsen and Kaufman, 1999; Kennedy et al., 2001b). This may be interpreted in two ways. Either the ⁸⁷Sr/⁸⁶Sr ratio of seawater changed little because the glaciation was short-lived (Jacobsen and Kaufman, 1999), contrary to paleomagnetic (Sohl et al., 1999) and subsidence (Hoffman et al., 1998a) data, or else the ratio was buffered by carbonate dissolution during the hypothesized snowball event and enhanced weathering in the greenhouse aftermath (Hoffman et al., 1998b). If atmospheric pCO₂ rose from a pre-glacial value of 0.0003 bar to 0.12 bar over the course of a snowball event (Caldeira and Kasting, 1992), then the Ca ion concentration of seawater would have to increase five-fold simply to maintain carbonate saturation (Fig. 10). This enormous Ca input would be accompanied by a flux of isotopically conservative Sr from carbonate dissolution and radiogenic Sr from silicate weathering. These inputs might greatly reduce or even eliminate the predicted lowering of ⁸⁷Sr/⁸⁶Sr during a snowball event of any given duration (Jacobsen and Kaufman, 1999). The large magnitude of the resulting seawater Sr reservoir, plus the rapid weathering of carbonates and young volcanics, would damp any increase in seawater ⁸⁷Sr/⁸⁶Sr due to silicate weathering, consistent with available data (Jacobsen and Kaufman, 1999; Kennedy et al., 2001b). Contrary to Kennedy et al. (2001b), the transient weathering flux is not limited to the timescale of the sea-level rise (<104 years) but continued long after as indicated by the negative δ¹³C anomaly (Fig. 9) which encompasses the entire transgressive-regressive post-glacial sequence. Accordingly, ⁸⁷Sr/⁸⁶Sr ratios in cap carbonates are broadly similar to pre-glacial values but should rise thereafter due to silicate weathering, as is observed (Shields et al., 1997).

Would eukoryotes survive?
Few biologists doubt that prokaryotic organisms could survive snowball events (Priscu et al., 1998; Gaidos et al., 1999; Vincent and Howard-Williams, 2000; Thomas and Dieckmann, 2002). Extant clades of eukaryotic algae are known from strata around 1.0 Ga (German, 1990; Butterfield et al., 1990, 1994; Butterfield and Rainbird, 1998; Butterfield, 2000) and they must have prevailed as well. A marked decline in eukaryotic and total microfloral diversity encompasses the late Neoproterozoic glacial events (Vidal and Knoll, 1982; Schopf, 1991; Knoll, 1994; Vidal and Moczydlowska-Vidal, 1997), and it is not until well after the Elatina/Nantuo glaciation (Table 1) that an abrupt increase in abundance, size, morphological complexity, and taxonomic diversity of acanthomorph acritarchs is observed (Zhang et al., 1998; Grey, 1998; Walter et al., 2000). The enigmatic, soft-bodied, Ediacara macrofossils—arguably dominated by cnidarian-grade metazoans—appears world-wide by 575-555 Ma (Jenkins, 1992; Fedonkin, 1992; Seilacher, 1992; Martin et al., 2000). In northwestern Canada, simple Ediacara-type fossils with radial symmetry occur in the immediate post-glacial sequences of both the Rapitan (Hofmann et al., 1990) and Ice Brook (Narbonne and Aitken, 1990; Narbonne, 1994) glaciations (Table 1). Many molecular-based age estimates of the last common ancestor of bilaterians predate the hypothesized snowball events (Runnegar, 2000), but there are good reasons to take these estimates lightly (Peterson and Takacs, 2001).

Assuming snowball events occurred, how did eukaryotic plankton survive? In what refugia did early metazoans, if they existed, ride out the climate shocks entering and exiting the hypothesized snowball events? Did their prolonged isolation contribute ultimately to evolutionary diversification? Potential refugia include brine channels in juvenile sea ice (Thomas and Dieckmann, 2002), leads and lanes in the tropical ice pack (Kirschvink, 1992; Lemke, 2001), tidal cracks along ice-grounding lines (Gaidos et al., 1999), transient meltwater ponds (Walker, 2001), thin equatorial sea ice (McKay, 2000), and hot springs around volcanic islands (Hoffman and Schrag, 2000). Snowball seawater would be laden with nutrients due to hydrothermal activity and limited organic productivity (Kirschvink, 2000). Algal “blooms’ and enhanced organic burial associated with rapid sedimentation when snowball oceans were uncorked would have released oxygen to the atmosphere (Kirschvink et al., 2000; Walter et al., 2000). This is supported by evidence of increased environmental oxygenation roughly coincident with both the Paleoproterozoic (Prasad and Roscoe, 1996; Rye and Holland, 1998; Farquhar et al., 2000; Murakami et al., 2001) and Neoproterozoic (Des Marais et al., 1992; Logan et al., 1995; Canfield and Teske, 1996) snowball eras (Fig 12). The isotopic signal of such organic-burial events might be muted if they were accompanied by a high carbonate burial flux, driven by weathering as we have argued. The rise of the Ediacaran biota has long been attributed to a rise in molecular oxygen (Runnegar, 1982; Knoll and Carroll, 1999), but it is not clear what stabilized pO₂ at higher levels following the perturbations.