Geological evidence indicates that grounded ice sheets reached sea level at all latitudes during two long-lived Cryogenian (58 and ≥5 My) glaciations. Combined uranium-lead and rhenium-osmium dating suggests that the older (Sturtian) glacial onset and both terminations were globally synchronous. Geochemical data imply that CO 2 was 10 2 PAL (present atmospheric level) at the younger termination, consistent with a global ice cover. Sturtian glaciation followed breakup of a tropical supercontinent, and its onset coincided with the equatorial emplacement of a large igneous province. Modeling shows that the small thermal inertia of a globally frozen surface reverses the annual mean tropical atmospheric circulation, producing an equatorial desert and net snow and frost accumulation elsewhere. Oceanic ice thickens, forming a sea glacier that flows gravitationally toward the equator, sustained by the hydrologic cycle and by basal freezing and melting. Tropical ice sheets flow faster as CO 2 rises but lose mass and become sensitive to orbital changes. Equatorial dust accumulation engenders supraglacial oligotrophic meltwater ecosystems, favorable for cyanobacteria and certain eukaryotes. Meltwater flushing through cracks enables organic burial and submarine deposition of airborne volcanic ash. The subglacial ocean is turbulent and well mixed, in response to geothermal heating and heat loss through the ice cover, increasing with latitude. Terminal carbonate deposits, unique to Cryogenian glaciations, are products of intense weathering and ocean stratification. Whole-ocean warming and collapsing peripheral bulges allow marine coastal flooding to continue long after ice-sheet disappearance. The evolutionary legacy of Snowball Earth is perceptible in fossils and living organisms.

Because of inconsistent usage in the current glacial sedimentology literature, we need to clarify that we use the terms “dynamic” and “stability” in reference to ice flow and ice extent, respectively. A “dynamic” glacier is one that flows, and an “unstable” one waxes and wanes. Finally, we do not use the words “hypothesis” and “theory” to distinguish degrees of confidence or certainty regarding Snowball Earth. Theory refers to the mechanism behind the phenomenon. The hypothesis is the postulate ( 24 ) that the phenomenon actually occurred in the geologic past.

To a geologist, it may seem that climate models are overly concerned with the Snowball state and with glacial states close to the Snowball bifurcation. Why not the more geologically interesting onsets and terminations ( Fig. 1 )? The reason for this is that current climate models used in deep-time investigations are meaningful only as equilibrium responses to prescribed forcings. The equilibration time is typically hundreds to thousands of model years. Snowball onsets and terminations are highly nonequilibrium situations on these time scales. These nonequilibrium states can be modeled but are computationally demanding.

We offer a few words for geologists who are more accustomed to the “inverse” approach. When we speak declaratively in describing the results of a climate model, there is no implication that nature necessarily behaved similarly. A model result is simply the equilibrium response of the model to a set of prescribed conditions. It is a job for geology to determine whether the record is consistent with model predictions. Like geology itself, models are a source of astonishment, wonder, and inspiration. Like geological observations, they are incomplete and subject to improvement. To help make the distinction clear, we will endeavor to use the present tense in describing model results and the past tense for inference based on geological data.

We begin by reviewing Cryogenian geochronology and paleogeography, followed by Snowball atmospheric dynamics and the hydrologic cycle, Snowball ice-sheet extent and variability, low-latitude sea ice-margin stability, sea-glacier dynamics, supraglacial cryoconite ecosystems, subglacial ocean dynamics, and “cap-carbonate” sequences unique to Cryogenian glacial terminations. We conclude with a summary of modeling results and their geological implications. Although we refer to them tangentially, we do not discuss the isotopic proxy records surrounding Cryogenian glaciation. To do so would easily double the length of the paper (and the list of authors).

This unpromising situation is now changing. Research on Snowball Earth climate dynamics has taken hold at leading institutions on four continents. Its motivation is not only to assist geology and geobiology but also to pursue potential applications for study of exoplanets, to gain insights from intermodel comparison, and to stimulate fresh perspectives on the Anthropocene. The goal of this paper is more modest. We survey progress in modeling the Snowball Earth atmosphere, cryosphere, hydrosphere, and lithosphere, specifically as it pertains to Cryogenian geology and geobiology. Such a review is timely because the recent development of a radiometric chronology for Cryogenian glaciation ( 63 ) has breathed new life into Neoproterozoic research. We build on an insightful pair of 6-year-old reviews of the same topic ( 115 , 116 ) and offer ours in the same ecumenical spirit.

Less clear is how the Cryogenian sedimentary and geobiological records should be interpreted within the context of Snowball Earth. The Cryogenian sedimentary record is widespread and accessible ( Fig. 4 ), fostering detailed studies in many areas ( 92 – 102 ). In contrast, the Cryogenian fossil record has low total and within-assemblage diversity ( 103 ). Molecular fossils ( 46 ) and macrofossils ( 104 – 107 ) are rare, and molecular “clock” estimates for early metazoan divergences ( Fig. 3 ) are imprecisely calibrated ( 108 ). Moreover, recycling of preglacial organic matter into glacial deposits is inevitable, because glaciers would have readily entrained organic-rich sediments on the continental shelves they traversed ( 109 ). Although both the sedimentary and fossil records have been cited as casting doubt on the Snowball Earth hypothesis ( 107 , 110 – 114 ), the reality is that, until recently, the concept itself was not fleshed out in sufficient detail to make reliable predictions regarding those records.

Self-termination occurs under circumstances that are falsifiable geologically. Snowball cryochrons should be long-lived, reflecting the enormous amount of CO 2 that must accumulate to counteract the high planetary albedo. Their onsets should be synchronous at low latitudes, and their terminations should be synchronous everywhere ( Fig. 1 ). Terminations should be accompanied by extraordinarily high atmospheric CO 2 levels, equivalent to 10 4 to 10 5 parts per million (ppm) by volume in the present atmosphere, giving rise to torrid greenhouse aftermaths as the surface darkens due to ice retreat ( Fig. 1 ). These predictions are increasingly supported by combined U-Pb and Re-Os geochronology ( 32 , 56 – 65 , 83 ) and by geochemical data ( 60 , 84 – 91 ), none of which existed when Cryogenian Snowball states were hypothesized ( 24 , 27 ).

Seafloor weathering represents another sink for atmospheric CO 2 on Snowball Earth ( 73 – 75 ). Cold bottom-water temperature slows the rate of seafloor weathering ( 76 ), but this is nullified by the long duration and low pelagic sediment flux of the Snowball ocean. Weathering rate increases when the pH of cold seawater falls below 7.0 in response to CO 2 rise ( 76 ). Current estimates of the CO 2 level required for Cryogenian Snowball termination range from 0.01 to 0.1 volume mixing ratio, dependent on the tropical ice albedo and other factors ( 77 – 82 ). Seafloor weathering, a sink for CO 2 and a source of alkalinity, is a potential factor in the longevity of the Sturtian cryochron ( Fig. 2A ).

Deposition of CO 2 ice at the poles in winter is a potential sink that might render a Snowball Earth irreversible ( 70 ). CO 2 ice, having a density of 1.5 g cm −3 , might sink gravitationally into the polar water ice, melting as it penetrates warmer ice at depth ( 70 ). On the other hand, CO 2 ice clouds warm the poles in winter by scattering outgoing radiation ( 71 , 72 ), yet dissipate in summer, allowing sunshine to reach the surface and sublimate CO 2 ice particles in the firn ( 70 ). As atmospheric CO 2 accumulates, the stability range for deposition of CO 2 ice initially grows because of higher saturation ( 70 ). However, an ice-covered planet with Earth-like obliquity and distance (1 astronomical unit) from a Sun-like star, with or without CO 2 ice clouds, remains far below the threshold for CO 2 ice deposition at any CO 2 level relevant to a Cryogenian Snowball Earth [see figure 2b and c in the study by Turbet et al. ( 70 )]. The risk of irreversibility due to CO 2 -ice deposition is not so easily dismissed for a Siderian (2.4 Ga) Snowball Earth ( Fig. 2B ), when the Sun was about 10% dimmer than in the Cryogenian.

The recognition that a panglacial state might be self-terminating ( 69 ), due to feedback in the geochemical carbon cycle, meant that its occurrence in the geological past could not be ruled out on grounds of irreversibility. “If a global glaciation were to occur, the rate of silicate weathering should fall very nearly to zero (due to the cessation of normal processes of precipitation, erosion, and runoff), and carbon dioxide should accumulate in the atmosphere at whatever rate it is released from volcanoes. Even the present rate of release would yield 1 bar of carbon dioxide in only 20 million years. The resultant large greenhouse effect should melt the ice cover in a geologically short period of time” [( 69 ), p. 9781]. Because Snowball Earth surface temperatures are below the freezing point of water everywhere, due to high planetary albedo, there is no rain to scrub CO 2 (insoluble in snow) from the atmosphere.

The Cryogenian period ( 50 ) encompasses the paired Neoproterozoic Snowball Earths and the brief nonglacial interlude ( Fig. 2A ). The term “cryochron” ( 27 ) was proposed for the panglacial epochs, on the assumption that their onsets and terminations were sharply defined in time, which now appears to be the case ( Table 1 ). The older cryochron has come to be known as “Sturtian” and the younger one has come to be known as “Marinoan,” after Sturt Gorge and Marino Rocks near Adelaide, South Australia, where they were recognized and mapped over 100 years ago ( 51 ). [As originally defined ( 52 ), these regional terms did not refer exclusively to the glacial epochs, but the original terminology has been superseded by the formal periods of the International Time Scale ( Fig. 2A ) ( 50 , 53 ). We find it convenient to follow the current informal international usage, wherein “Sturtian” and “Marinoan” identify the cryochron and its immediate aftermath, including the respective postglacial cap-carbonate sequences.] The nonglacial interlude separating the cryochrons was 9 to 19 million years (My) in duration ( Fig. 2A ).

( A ) Black bands indicate durations of the Sturtian and Marinoan cryochrons ( Table 1 ). The graded start to the Marinoan cryochron denotes chronometric uncertainty, not gradual onset. Ellipse F-LIP shows the possible age span of the Franklin large igneous province (LIP) ( 32 , 127 ). ( B ) Snowball Earth chrons (black), regional-scale ice ages (medium gray), and nonglacial intervals (light gray) since 3.0 Ga. Ellipse GOE is centered on the Great Oxidation Event, as recorded by the disappearance of mass-independent S isotope fractionations ≥0.3 per mil (‰) in sedimentary sulfide and sulfate minerals ( 484 ). The dashed gray line indicates questionable glaciation.

Ice-line latitude as a function of solar or CO 2 radiative forcing in a one-dimensional (1D) (meridional) energy-balance model of the Budyko-Sellers type ( 3 , 4 ), showing three stable branches (red, green, and blue solid lines) and the unstable regime (dashed line). Yellow dots are stable climates possible with present-day forcing. Black arrows indicate nonequilibrium transitions. In response to lower forcing, ice line migrates equatorward to the ice-albedo instability threshold (a), whereupon the ice line advances uncontrollably to the equator (Eq) (b). With reduced sinks for carbon, normal volcanic outgassing drives atmospheric CO 2 higher over millions to tens of millions of years ( 73 ) until it reaches the deglaciation threshold (c). Once the tropical ocean begins to open, ice-albedo feedback drives the ice line rapidly poleward (in ~2 ky) ( 327 ) to (d), where high CO 2 combined with low surface albedo creates a torrid greenhouse climate. Intense silicate weathering and carbon burial lower atmospheric CO 2 (in 10 7 years) ( 164 ) to (e), the threshold for the reestablishment of a polar ice cap. The hysteresis loop predicts that Snowball glaciations were long-lived (b and c), began synchronously at low latitudes (a and b), and ended synchronously at all latitudes under extreme CO 2 radiative forcing (c and d). The ocean is predicted to undergo severe acidification and deacidification in response to the CO 2 hysteresis. Qualitatively similar hysteresis is found in 3D general circulation models (GCMs). Pco 2 , partial pressure of CO 2 ; wrt, with respect to.

The synchroneity of Franklin LIP emplacement with the onset of low-latitude glaciation ( 32 ) raises the possibility that volcanism was the proximal trigger for the Sturtian cryochron ( 200 ). Historical flood-basalt fissure eruptions generated thermal plumes that injected sulfate aerosol precursor gases (SO 2 and H 2 S) into the stratosphere intermittently for months ( 201 ), and more voluminous LIP eruptions likely extended this time scale to years ( 202 ). Many Franklin LIP sills and dykes are highly enriched in S, 10 2 to 10 5 ppm ( 203 ), because of contamination by sulfate evaporites they intrude within the pre-Sturtian Neoproterozoic Shaler Supergroup ( 204 ). Sulfur chemistry, aerosol microphysics, and radiative energy-balance modeling ( 200 ) reveal the importance of the ambient climate and the paleolatitude of eruption. A cold pre-Sturtian climate, related to the Rodinia breakup, was vulnerable to stratospheric aerosol forcing because the tropopause was lower in elevation and the ocean mixed layer was closer to the freezing point ( 200 ). Sulfate aerosol increases the planetary albedo, and the resultant cooling effect is maximized if injection occurs at low latitudes, where insolation is strongest ( 200 ). Accordingly, the Franklin LIP had unique climatic consequences because S-rich lavas were erupted at high rates at the paleoequator in an already-cold climate. Concurrent CO 2 emissions were small, relative to an ambient atmospheric reservoir of ≤3000 ppm ( 200 ), and were compensated by enhanced weathering on the time scale of LIP emplacement.

The Rodinia breakup was also associated with the emplacement of basaltic LIPs at ~825 Ma (Guibei-Willouran LIP), ~800 Ma (Suxiong-Xiaofeng LIP), ~775 Ma (Gunbarrel-Kinding LIP), ~755 Ma (Mundine Well LIP), and ~717 Ma (Franklin LIP). The last one was emplaced squarely across the paleoequator ( 34 , 35 , 127 , 191 ). Weathering of basaltic rock consumes CO 2 more rapidly than weathering average upper continental crust [( 192 , 193 ); but see the study by Jacobson et al. ( 194 )], and basalt weathers most rapidly at high temperatures ( 195 ). Basaltic rocks including pre-Sturtian LIPs are strongly enriched in phosphorus relative to average upper crust, possibly leading to enhanced organic productivity and burial ( 193 , 196 ), depending on the controls of apatite dissolution kinetics (that is, grain size and soil-water pH). Elevated rates of fractional organic burial are indicated by enriched C isotopes (δ 13 C ≥ 5‰ Vienna Pee Dee belemnite) in marine carbonate deposited from ~825 to 717 Ma [( 174 , 197 ); but see the study by Shields and Mills ( 198 )]. Enhanced weathering of young basaltic rocks is supported by relatively nonradiogenic Sr, Os, and Nd isotopic compositions observed in pre-Sturtian carbonate, organic matter, and shale, respectively ( 60 , 174 , 197 ). Enhanced CO 2 consumption and global cooling due to the breakup of the Rodinia supercontinent and the emplacement of a temporal cluster of LIPs ( 174 , 186 , 196 , 199 ) purportedly moved the climate system closer to the Snowball bifurcation ( Fig. 1 ).

The change in paleogeography over the Cryogenian period ( Fig. 5 ) reflects the breakup of the Rodinia supercontinent at this time. The breakup presumably reduced CO 2 greenhouse warming in two ways. First, moistening of previously arid lands in the supercontinental interior adjusted CO 2 to a lower level because of increased weathering efficiency ( 186 ), defined as the silicate weathering rate globally at any given CO 2 level. A less marked cooling followed the breakup of Pangea ( 187 ), perhaps because an equatorial ocean (Tethys) already existed within the Pangea supercontinent and because of the low albedo of tropical forests, nonexistent during the Cryogenian period. Second, the total length of all continental margins rose, resulting in higher rates of global sediment accumulation and consequently organic burial ( 188 , 189 ). The apparent absence of polar continents in Cryogenian paleogeography ( Fig. 5 ) weakened the protection against uncontrolled cooling normally afforded by silicate-weathering feedback. If global weathering rate declines because of the glaciation of polar continents, then atmospheric CO 2 stabilizes or even rises. When polar continents are absent, this protective feedback vanishes ( 190 ). The ice-albedo feedback is then driven by oceanic ice (that is, sea ice plus shelf ice) alone, and the Snowball bifurcation ( Fig. 1 , point a) can be reached with little prior development of grounded ice sheets ( 18 ). This may explain why regional-scale ice ages ( Fig. 2 ) have not been observed in Cryogenian time.

The tropical bias in Cryogenian continentality is postulated to have lowered global surface temperatures by raising the planetary albedo ( 24 ). Experiments with a coupled atmosphere-ocean GCM, ECHAM5/MPI-OM, support this supposition, yielding a global mean surface temperature with Cryogenian paleogeography ~3°C lower than present, because of both higher albedo and suppressed water-vapor greenhouse ( 18 ). However, the problem is not a simple one—Tropical continents also suppress evaporative cooling ( 182 ), and an absence of high-latitude continents reduces summer snow cover, leading to warmer high-latitude oceans and less sea ice ( 183 ). Experiments with an atmospheric GCM, CAM3.1, coupled to a mixed-layer ocean model and a land model, CLM (community land model), yield a warmer-than-present climate with Cryogenian paleogeography because of decreased tropical cloud cover over land and intensification of the Walker circulation, leading to more (rather than less) water-vapor greenhouse ( 184 ). However, on the time scale of the geochemical cycle of carbon, more continental area in the deep wet tropics should yield a globally colder climate because of enhanced silicate weathering ( 185 ).

In the absence of vascular plants, continental surface albedos were presumably higher than present-day albedos in the tropical and mid-latitude wet zones, but transpiration and hence evaporative cooling were less. Silicate weathering and silicate-weathering feedback, the ultimate climate thermostat ( 69 ), may have been weaker in the absence of rooted plants ( 22 ), which increase soil acidity by CO 2 pumping and organic acid production. Rooted plants also stabilize hillslopes, which lengthens soil residence time in the weathering zone, but retards the entry of fresh rock into that zone. Higher atmospheric CO 2 , adjusted to the 6 to 7% dimmer Cryogenian Sun ( 69 ), would have lowered soil-water pH, offsetting the absence of rooted plants to some degree. Lower seawater pH would have enhanced seafloor weathering at low temperatures ( 172 , 173 ), but the enlarged ocean-atmosphere C reservoir would have weakened the silicate-weathering feedback because of the climate’s logarithmic dependence on CO 2 .

Cryogenian glacial deposits can now be retrolocated on paleogeographic maps ( Fig. 5 ), constructed by means of paleomagnetically constrained interpolation between an inferred configuration of the Rodinia supercontinent at 780 Ma and that of the Gondwana megacontinent at 520 Ma ( 34 , 179 , 180 ). The main uncertainty is Rodinia: Adopt a different configuration for Rodinia ( 181 ) and the specifics of Cryogenian paleogeography will be quite different, even if the generalities remain. Notable among the generalities are the absence of polar continents and the large low-latitude land area ( Fig. 5 ).

Cryogenian glacial deposits ( 28 , 175 ) are reliably documented in >90 formations on 22 paleocontinents and microcontinents worldwide ( Fig. 4 ). Areas where they are absent, such as northeastern North America, simply lack Cryogenian strata and should not be considered to have been ice-free. Certain Sturtian and Marinoan deposits were formed by glaciers that were grounded below sea level at low paleomagnetic latitudes ( 28 , 31 , 33 , 34 ). Glacial deposits of both cryochrons are bounded regionally by thick, shallow-water, nonskeletal, carbonate sequences ( 24 , 25 , 34 , 143 , 144 , 176 ). In both warm and cool global climates, nonskeletal carbonate production occurs preferentially in the warmest parts of the surface ocean ( 177 , 178 ), reflecting the temperature and inverse pressure dependence of carbonate saturation state. Moreover, broad carbonate platforms lack mountains from which glaciers could descend. This was the logic that justified the Snowball Earth hypothesis before proof of synchroneity was available. If the warmest surface areas were glaciated, then colder areas must also have been frozen.

Sinks for CO 2 on Snowball Earth include subglacial weathering of continental crust ( 171 ), including LIPs, low-temperature alteration of young oceanic crust on the flanks of seafloor spreading ridges ( 75 , 76 , 172 , 173 ), and organic burial (see the “Meltwater flushing, the carbon cycle, and atmospheric oxygen” section). Geochronologic and paleomagnetic data demonstrate that the windward margin of Rodinia in the deep tropics was resurfaced by flood basalt of the Franklin LIP ( Fig. 5 ) just before the Sturtian glaciation ( 32 ). Nonradiogenic Sr, Os, and Nd isotope compositions in sediments show that LIP and/or seafloor weathering had been dominant for the preceding 15 My ( 60 , 174 ). It has been suggested that the removal of volcanic plateaus by Sturtian glacial erosion rendered continents less reactive, thereby accounting for the shorter duration of the Marinoan glaciation ( 174 ). A potential problem for this idea is that debris from the ca. 825-Ma Wooltana (LIP) volcanics in South Australia is prominent in both the Sturtian and Marinoan glacial deposits in that region, although isotopic tracers globally become more radiogenic after the Sturtian glaciation, consistent with progressive LIP removal ( 174 ).

Comparison of stratigraphic thickness of Marinoan and Sturtian cryochrons (blue dots and whiskers: mean ± 1σ for 492 records) with Phanerozoic shallow glaciomarine accumulation (red dots: 6733 records) and nonglacigenic terrigenous shelf accumulation (gray: ±1σ band for 32,892 records), plotted by duration of deposition. The yellow diamond represents the 580-Ma Gaskiers glaciation ( 247 ). Comparison of comparable durations is mandated because accumulation rate (dashed contours in meters per million years) decreases as averaging time increases, due to stratigraphic incompleteness ( 240 ). Data are from Partin and Sadler ( 169 ) and Sadler and Jerolmack ( 242 ). Cryogenian glacial deposits accumulated 3 to 10 times more slowly than younger glacial deposits of comparable duration ( 169 ).

If rift-related paleotopography was greater during the Sturtian cryochron, then erosion and sedimentation rates should have been greater as well. This prediction is not supported by data on their respective sediment accumulation rates ( Fig. 6 ) ( 169 ). Sturtian sections have median and average thicknesses that are two and four times greater, respectively, than those of Marinoan sections. However, the average accumulation rate is actually lower for Sturtian than for Marinoan sections because of the disproportionate averaging times ( Fig. 6 ) ( 169 ). The apparent contradiction could stem from a subtlety. Whereas it is the extent of ice that governs Snowball albedo and hence the CO 2 threshold for deglaciation, it is the flux of ice (at equilibrium lines) that appears to govern rates of erosion and sedimentation by active glaciers ( 170 ). More rugged Sturtian landscapes with more ice and less dust could be consistent with sedimentation rate data ( Fig. 6 ) if Sturtian ice sheets were, on average, less dynamic. For example, modeling indicates that ice sheets are largest on Snowball Earth when CO 2 is low and the ice flux is weak ( 100 ). For this to occur, low surface temperatures and thus a feeble hydrologic cycle would need to be maintained ( 81 ) during much of the Sturtian cryochron. This scenario is unlikely: The most-rapid warming should occur at the start of a cryochron because of its logarithmic dependence on CO 2 and because of the pH dependence at low temperatures of CO 2 consumption by seafloor weathering ( 76 ).

Global paleogeographic reconstructions in Mollweide projection for ( A ) Marinoan termination at 635 Ma and ( B ) Sturtian onset at 720 Ma ( 34 ). Red lines are oceanic spreading ridge-transform systems, and dark blue lines with barbs are inferred subduction zones. Stars are glacial-periglacial formations ( Fig. 4 ), red stars are formations with synglacial iron formation, and green stars indicate occurrences of authigenic and/or seafloor barite in Marinoan postglacial cap dolostone ( Fig. 4 ). Cryogenian glaciation was coeval with the breakup of supercontinent Rodinia. Paleocontinent Laurentia (Laur) is fixed in latitude and declination at 720 Ma by paleomagnetic data (n = 87 sites) from the Franklin LIP (purple dashed line) in Arctic Canada ( 32 , 33 , 127 ). Paleocontinents South Australia (SA) and South China (SCh) are similarly fixed at 635 Ma by paleomagnetic data from the Nuccaleena Formation cap dolostone ( 33 ) and the Nantuo Formation glacial diamictite ( 371 ), respectively. Other paleocontinents are Amazonia (Amaz), Avalonia (Av), Baltica (Balt), Cadomia (Ca), Congo, Dzabkhan (Dz) in Mongolia, East Antarctica (EAnt), East Svalbard (ESv), India (Ind), Kalahari (Kal), Kazakhstan (Kaz), North Australia (NA), North China (NCh), Oman (Om), Rio de la Plata (P), São Francisco (SF), Siberia (Sib), Tarim (Tm), and West Africa (WAfr). The paleolocation of Oman is uncertain; alternatively, it could restore west of India. The Sturtian location of South China opposite Laurentia is controversial; it might have been closer to its Marinoan position ( 368 ).

The difference in cryochron duration should relate to the low-latitude albedo or to sinks for CO 2 that could retard its rise in the atmosphere. In many areas, geological observations imply that Sturtian glacial deposits accumulated during the active rifting of Rodinia ( Fig. 5 ), whereas Marinoan ones accumulated during the drift stage of post-rift subsidence. If true, active rift flanks may have increased Sturtian paleotopography, resulting in more complete coverage of the continents by ice sheets ( 165 , 166 ), most importantly by glacial flow into the equatorial zone of sublimation. This would have raised the planetary albedo. Increasing the ice coverage would also reduce the atmospheric dust load, allowing long-wavelength radiation to escape more readily ( 77 , 115 ). A lower dust flux would additionally raise albedo in the sublimation zones of glaciers ( 78 , 79 , 167 , 168 ). More ice and less dust means that more CO 2 must accumulate to deglaciate ( 116 ).

Another surprising aspect of the new chronology is the brevity of the nonglacial interlude between the cryochrons. When the Marinoan began, a Snowball Earth had terminated less than 20 My earlier. When the Sturtian began, no low-latitude glaciation had occurred for 1.7 Gy ( Fig. 2B ). Might this contrast in circumstance relate to the inequality of the cryochrons? Might it rationalize the geochemical distinctions between them ( 130 , 131 , 158 )? Nearly all synglacial Cryogenian iron formations ( Fig. 4 ) are Sturtian in age ( 159 , 160 ), whereas all known occurrences of barite (BaSO 4 ) in cap dolostones are Marinoan ( 91 , 155 , 161 – 163 ). Nearly all the isotopic evidence for anomalous atmospheric CO 2 comes from the shorter Marinoan cryochron ( 86 ) or its aftermath ( 84 , 85 , 87 – 91 , 162 , 163 ). This reflects the absence in the Sturtian of glaciolacustrine calcite and postglacial barite and the incomplete development of Sturtian cap-carbonate sequences, discussed later. There is no reason to assume that Sturtian deglaciation required less CO 2 than the Marinoan. The 10- to 20-My nonglacial interlude is of the same duration as the estimated time scale for the post-Snowball drawdown of atmospheric CO 2 ( 164 ).

The most striking aspect of the new Cryogenian chronology is the grossly unequal duration of the cryochrons ( Fig. 2A ). The Sturtian lasted 4 to 19 times longer than the Marinoan. This could be the result of a small difference in equatorial surface albedo, because surface warming has a logarithmic dependence on CO 2 concentration. Because CO 2 is assumed to have accumulated linearly or at a decreasing rate over time during a cryochron ( 73 ), subject to sequestering as CO 2 ice or consumption by seafloor weathering, as mentioned earlier, even a small increase in the radiative threshold for melting could require a long time to achieve because of the high CO 2 concentration required to trigger deglaciation ( Fig. 1 ).

The termination of the Marinoan cryochron is globally marked by a laterally continuous unit of pale, pinkish-to-beige–colored, peloidal dolostone with wave-generated sedimentary structures ( 97 , 130 , 154 ). This so-called “cap dolostone” sharply overlies the last glacial and periglacial deposits without evident hiatus. It was deposited during glacioeustatic flooding resulting from global deglaciation ( 154 – 156 ). The date of Marinoan deglaciation is tightly constrained between 636.0 and 634.7 Ma by U-Pb (CA-ID-TIMS) zircon ages from volcanic ash layers within the youngest glacial deposits in Namibia and Australia and at the top of the cap dolostone in South China ( Table 1 ). Because of the uncertainty in the date of its onset, the duration of the Marinoan cryochron is loosely constrained between 3.0 and 15.2 My. A thermal subsidence model of the carbonate passive margin (upper Otavi Group) in northern Namibia ( 157 ) implies ~6 My for the Marinoan cryochron together with its postglacial depositional sequence (Maieberg Formation), in which the cap dolostone forms the base.

The nonglacial interlude between the Sturtian termination and the Marinoan onset had a total duration between 8.6 and 20.3 Ma ( Table 1 ). The low or high ends of this range agree with an astrochronological estimate of 6 to 8 My for the same interlude in East Svalbard, assuming that 0.5-m-scale rhythms of dolomitic siltstone in shale were forced by orbital precession ( 138 ) or obliquity, respectively. However, whether the rhythms are truly periodic remains to be seen. Comparable in duration to the shorter Phanerozoic epochs (for example, Oligocene, Middle Jurassic, and Middle Devonian), the nonglacial interlude can be globally correlated. In East Greenland ( 139 ), East Svalbard ( 138 ), and South China ( 140 – 142 ), 0.2 to 0.3 km (uncompacted) of fine-grained terrigenous strata appear. In western Mongolia ( 143 ) and northern Namibia ( 144 ), one finds 0.6 km of limestone and dolostone. In Idaho ( 133 ) and Oman ( 145 ), there are 0.7 to 0.8 km of fine- and coarse-grained terrigenous deposits. Northwest Canada (Mackenzie Mountains) ( 146 – 149 ) and South Australia ( 150 – 153 ) have 1.3 and 3.0 km, respectively, of mixed terrigenous and carbonate strata. Accommodation created during the Sturtian glaciation and high rates of weathering and erosion in the Sturtian aftermath account for the high sedimentation rates (<0.1 km My −1 ). The absence of large ice sheets is inferred from peritidal cycles in platform carbonates that lack karstic disconformities and other evidence of forced regressions [( 144 , 152 ); but see the study by Day et al. ( 149 )].

The onset of the Marinoan cryochron is only loosely constrained between 649.9 and 639.0 Ma by a pre-Marinoan Re-Os isochron age of 645.1 ± 4.8 Ma ( 137 ) for organic-rich shale (Tindelpina Member, lower Tapley Hill Formation) in South Australia and a syn-Marinoan U-Pb (CA-ID-TIMS) zircon age of 639.3 ± 0.3 Ma ( 83 ) from a volcanic tuff within glaciomarine diamictite (Ghaub Formation) in northern Namibia. Additional constraints on the Marinoan onset come from U-Pb (SIMS) zircon dates from tuffs within pre-Marinoan strata (Datangpo Formation) in South China ( Table 1 ).

Some authors infer that the Sturtian cryochron encompassed multiple discrete glaciations ( 132 – 135 ). There is sedimentological evidence for multiple advances and retreats of Sturtian ice-sheet margins and ice grounding lines ( 92 , 95 , 136 ). However, the cap limestone is unique to the final ice retreat, and wherever it has been dated, its age is ~659 Ma ( Table 1 ). Nowhere have multiple cap carbonates been found within either cryochron, implying that the terminal deglaciations were unique events.

The termination of the Sturtian cryochron is globally marked by an organic-rich limestone (hereafter “cap limestone”) that sharply but conformably overlies glacial or periglacial deposits ( 130 , 131 ). Re-Os isochron ages from cap limestones in Australia, Laurentia (Northwest Territories, Canada), and Mongolia, and a U-Pb (SIMS) zircon date from a volcanic ash layer within terminal glaciomarine deposits in South Australia, tightly constrain the date of the Sturtian termination between 659.3 and 658.5 Ma ( Table 1 ). Accordingly, the duration of the Sturtian cryochron was between 57.0 and 59.0 My, nearly as long as the entire Cenozoic era.

The onset of the Sturtian cryochron at low (21° ± 3°N) paleolatitude is tightly constrained between 717.5 and 716.3 Ma by U-Pb (CA-ID-TIMS) zircon ages ( 32 ) from volcanic rocks in central Yukon (Canada). Zircon ages from South China and Arctic Alaska are consistent with these constraints ( Table 1 ). The Sturtian onset is indistinguishable in time from the most reliable U-Pb baddeleyite ages of 716.33 ± 0.54 Ma ( 32 ) and 716 ± 1 Ma ( 127 ) from mafic sills and dikes in different parts of the Franklin LIP of northern Laurentia. Franklin dikes, sills, and lavas extend across northern Canada from Alaska to Greenland ( 127 ) and Siberia ( 128 ), implying that the original volcanic plateau may have covered an area of >3 × 10 6 km 2 , making it among the largest terrestrial LIPs of all time ( 129 ).

Re-Os geochronology is a more recent development, pioneered less than 30 years ago ( 123 ) and successfully applied to a Paleozoic black shale in 2002 ( 124 , 125 ). Early attempts to date Neoproterozoic shales had mixed success, but after an improved extraction technique was introduced ( 126 ), intended to screen out nonhydrogenous (that is, detrital) Re and Os, isochron ages for Neoproterozoic organic-rich shales and limestones became generally compatible with existing U-Pb dating [( 63 and Table 1 )]. As with all isochron methods, sampling and sample selection are critical.

The time scale for Cryogenian glaciation ( Table 1 ) is based on U-Pb geochronology of the mineral zircon (ZrSiO 4 ), separated from volcanic rocks including far-traveled ash layers (tuffs), and on Re-Os isochron ages for the deposition of organic-rich sediments. The former is the most precise and most accurate dating method in deep-time applications because of the refractory nature of zircon, its low nonradiogenic Pb content, and an internal “concordance” test between two independent decay schemes ( 238 U→ 206 Pb and 235 U→ 207 Pb) having different decay constants. The long history, current methodologies, and error-propagation procedures of U-Pb geochronology have recently been reviewed ( 117 – 121 ). An important recent development was chemical abrasion of zircon grains ( 122 ), which improves concordance through selective removal of damaged grain fractions that are susceptible to radiogenic Pb loss. Chemical abrasion coupled with ID-TIMS (CA-ID-TIMS) is the most accurate dating method. Early Pb loss is an insidious problem, resulting in ages that are too young but still concordant. Zircon grains having high U contents are most susceptible to early Pb loss and are routinely avoided. Basaltic rocks typically lack zircon but may contain the mineral baddeleyite (ZrO 2 ), which can also be dated by U-Pb but with technical issues that are less well resolved than for zircon.

Before 2004 ( 56 ), no Cryogenian glacial deposit had been directly dated radiometrically. The Snowball Earth hypothesis ( 24 ) was then 12 years old, and the prospect of testing it geochronologically appeared remote. Today, the onset of the Sturtian cryochron and the terminations of both cryochrons are tightly constrained at a million-year resolution by U-Pb and Re-Os geochronology on multiple paleocontinents ( Table 1 ). This is a remarkable achievement and, more than any other factor, encourages the hypothesis to be taken seriously by geologists.

The observed contrast between Cryogenian and younger glacial accumulation rates ( Fig. 6 ) is striking ( 169 ). On average, Cryogenian accumulation rates were 3 to 10 times slower than for younger glaciations of comparable duration. This is remarkable given that the Phanerozoic data with long averaging times come exclusively from polar glaciations, Paleozoic Gondwana and Cenozoic Antarctica, whereas the Cryogenian data come from low to mid-paleolatitudes ( 169 ). Representing a database of 242 Sturtian and 326 Marinoan measured sections—incomplete, poorly exposed, and zero-thickness sections excluded—the anomalous Cryogenian accumulation rates are most readily explained by the attenuated hydrologic cycle of the “hard” Snowball Earth ( 169 ). In contrast, the sediment accumulation rate averaged over the short-lived (<350 ky), mid-latitude, 580-Ma Gaskiers glaciation ( 247 ) is indistinguishable from that of Cenozoic glaciations ( Fig. 6 ). The Cryogenian data ( 169 ) also allow for the possibility (to be discussed in the next section) that low-latitude ice sheets on Snowball Earth may have receded well before the cryochron-ending deglaciation of the ocean ( 100 ).

Because erosion and sedimentation rates appear to scale with the ice flux at the equilibrium line in active glaciers ( 170 ), we expect these rates to increase over the course of a Snowball cryochron ( Figs. 8 and 9 ). However, the average rates over entire cryochrons are expected to be smaller than for regional-scale glaciations because of the relatively weak hydrological cycle ( 81 ). With the development of a Cryogenian radiometric chronology ( Table 1 ), average sediment accumulation rates over entire cryochrons can now be calculated because the synglacial strata are generally well defined stratigraphically ( 175 ). Average accumulation rates for all sediments decline as averaging times increase ( Fig. 6 ) because of stratigraphic incompleteness ( 240 – 242 ). Therefore, Cryogenian accumulation rates should only be compared with younger glacial regimes of comparable duration ( Fig. 6 ) ( 169 ). The East Antarctic Ice Sheet, for example, has been in continuous existence for 14 My ( 243 , 244 ), which is close to the maximum duration of the Marinoan cryochron. Late Paleozoic glaciation of polar Gondwana caused the global mean sea level to fluctuate by ≥40 m on orbital time scales from late Viséan through end-Kungurian time ( 245 ), an interval of ~65 My that exceeds the Sturtian cryochron in duration. Smaller sea-level fluctuations are inferred at times within this interval ( 246 ), but it is unclear whether they imply less ice or smaller orbital-scale fluctuations.

At low CO 2 (0.1 mbar), ice-equivalent accumulation rates outside the tropics in the (non-FOAM) GCMs ( Fig. 9E ) are on the order of 1.0 mm year −1 (~1.0 km My −1 ), whereas rates in the inner subtropics reach a few centimeters per year or a few kilometers every 100 ky. These rates increase around fivefold when CO 2 is raised to 100 mbar ( Fig. 9F ). The rates pertain to sea level (sea ice) and cannot be directly extrapolated to the elevated surfaces of grounded ice sheets. Early studies of ice-sheet development in atmospheric GCMs with Cryogenian paleogeography suggested that tropical ice sheets would thicken and achieve a dynamic steady state within a few 100 ky after the tropical ocean froze over ( 165 , 239 ). This implies that glacial erosion and sedimentation occur continuously during most of the time span of a Snowball cryochron. However, this conclusion was tentative because the simulated air circulation and precipitation patterns in the early studies ( 165 , 239 ) were not iteratively coupled to ice-sheet topographic development. New modeling with improved coupling will be described in the “Ice-sheet stability and extent under precession-like forcing and variable CO 2 ” section.

The Snowball atmosphere is dry because it is cold, not because sources of water vapor are absent. The saturation water-vapor pressure is exponentially dependent on temperature (Clausius-Clapeyron relation) but is not much lower over ice than over supercooled water of equal temperature. The latent heats of evaporation and sublimation (direct conversion of ice to water vapor) differ by only 13.2%. Sublimation of ice is therefore a viable source of water vapor. Because of the “reversed” Hadley circulation, the main source of water vapor on Snowball Earth is equatorial sea ice, and the zones of highest net accumulation are the inner subtropics ( Fig. 9 , E and F). Narrow secondary bands of net sublimation occur at ~25° latitude, possibly related to export of moist air by mid-latitude eddies ( 235 ). Beyond that latitude, a low rate of net accumulation occurs everywhere, and the polar deserts of the modern climate are absent ( 81 ).

One of the GCMs, FOAM, is an outlier, yielding tropical surface temperatures 7 to 11 K colder than the other models ( Fig. 8 , A and B), equivalent to changing CO 2 by a factor of 10 to 100 ( 80 – 82 ). FOAM remains far below the deglaciation point at any geologically feasible CO 2 level, even with prescribed albedos of 0.5 and 0.75 (broadband average) for ablative and snow-covered ice surfaces, respectively ( 227 ). The reason is that FOAM underpredicts absorption of outgoing long-wavelength radiation by cloud condensate, for example, ice particles ( 80 – 82 , 116 ). This was originally considered to be unimportant because the dry Snowball atmosphere was assumed to lack optically thick clouds ( 227 , 228 ). However, a cloud-resolving model (System for Atmospheric Modeling, version 6.10.4) produces ≥10 W m −2 of cloud radiative forcing under a wide range of microphysical parameters ( 82 ), consistent with the non-FOAM GCM results ( Figs. 8 and 9 ) using the same uniform 0.6 surface albedo. This implies that Snowball deglaciation is not problematic from a modeling perspective, given the geological constraints on cryochron duration ( Table 1 ) and CO 2 concentration ( 85 – 91 ).

The zone of ascending air is closely tied to the zone of maximum insolation because of the low thermal inertia of the solid surface. When the ascending branch moves off the equator during the seasonal cycle, the equator falls under the descending branch of the Hadley cell ( Fig. 9 , A and B). Because the ascending branch spends more time away from the equator than above it, the annual mean Hadley circulation ( Fig. 9 , C and D) is characterized by equatorial downwelling and subtropical upwelling ( 81 , 116 ). This “reverse” Hadley circulation is unlike any (low-obliquity) climate with tropical surface water ( 236 , 237 ) and is more like the present climate of Mars ( 238 ) but with a denser atmosphere.

The Snowball atmosphere ( 80 – 82 , 116 , 226 – 229 ) is cold, due to high planetary albedo, and therefore holds little moisture. Latent heat contributes little to convection. The solstitial height of the equatorial tropopause is only 6 to 8 km above sea level ( 81 ), ~10 km less than in the present (geologically cold) climate. Outside the tropics, circulation is weak and dominated by subsidence in winter ( Fig. 9 , A to D). The Hadley circulation, although shallower than in the modern tropics, is actually significantly stronger ( Fig. 9 , A to D) ( 80 – 82 , 228 , 229 , 235 ). Vertical transport of moisture is much smaller than in the modern tropics, because the air is dry. Consequently, the ascending branch must be stronger to move enough heat to balance radiative cooling aloft ( 82 ).

Because the solid surface has little thermal inertia, surface temperatures closely track the instantaneous insolation forcing ( 81 , 116 , 226 – 229 ). Diurnal and seasonal cycles are amplified. At solstice ( Fig. 8B ), temperatures are uniformly warm (relatively speaking) in the summer hemisphere, whereas the winter hemisphere becomes extremely cold. The seasonal temperature change is large at all latitudes, consistent with geological evidence such as periglacial sand-wedge structures at low paleomagnetic latitudes ( 230 – 233 ), obviating the need for high orbital obliquity ( 231 , 233 ). Temperature inversions develop in the winter hemisphere ( 81 , 116 , 226 , 228 ), decoupling the surface from winds and from the greenhouse effect ( 81 ), which is set by the radiative balance at the top of the atmosphere ( 234 ).

See Fig. 8 caption for the prescribed conditions. January mean mass Eulerian stream function for models SP-CAM, LMDz, and ECHAM at ( A ) CO 2 = 0.1 mbar and ( B ) CO 2 = 100 mbar. σ is air pressure as a fraction of the surface air pressure. Clockwise atmospheric circulation is depicted by thin solid lines, counterclockwise circulation is depicted by thin dashed lines, and the zero stream function is depicted by thick solid lines. Contour interval is 50 × 10 9 kg s −1 . Maximum mass stream functions in January increase by a factor of ~1.5 between 0.1 and 100 mbar of CO 2 . ECHAM is unstable at CO 2 = 100 mbar. Note the ascending flow in January between 10°S and 30°S and the descending flow between 10°S and 20°N. Red shade indicates regions where the local Rossby number is greater than 0.5, meaning that inertial forces dominate over Coriolis forces. ( C and D ) Same as (A) and (B) but depict the annual mean mass Eulerian stream function. Contour interval is 20 × 10 9 kg s −1 . Note that air descends in the inner tropics in the annual mean and ascends in the subtropics. This governs the surface hydrologic balance (E and F). Annual and zonal mean precipitation minus sublimation with ( E ) CO 2 = 0.1 mbar and ( F ) CO 2 = 100 mbar. Note the net sublimation in the inner tropics and the net accumulation in the near subtropics, opposite to the present climate. The hydrologic cycle amplifies by nearly a factor of 10 between 0.1 and 100 mbar of CO 2 but is muted in FOAM due to cold surface temperatures ( Fig. 8 , A and B). LMDz exhibits grid-scale noise in (F).

Colors assigned each GCM as indicated in (B). Solid lines, CO 2 = 0.1 mbar; dashed lines, CO 2 = 100 mbar. To isolate differences in atmospheric behavior among models, surface albedo was set to 0.6 everywhere, eliminating differences between ablative and snow-covered ice. Topography and the radiative effect of aerosols were set to zero, as were all greenhouse gases other than CO 2 and H 2 O. The solar constant was set to 94% (present-day value; 1285 W m −2 ), obliquity was set to 23.5°, and eccentricity was set to zero. ( A ) Annual and zonal mean surface temperature. ( B ) January zonal mean surface temperature. ( C ) Sea-glacier thickness in meters (increasing downward to simulate depth below the ice surface). ( D ) Meridional (equatorward) ice velocity in meters per year. Models used are as follows: CAM (Community Atmosphere Model) ( 489 ), SP-CAM (Super-Parameterized Community Atmosphere Model) ( 490 , 491 ), FOAM (Fast Ocean Atmosphere Model) ( 492 ), ECHAM (European Centre Hamburg Model) ( 493 ), LMDz (Laboratoire Météorologie Dynamique Zoom) ( 256 ), and GENESIS (Global Environmental and Ecological Simulation of Interactive Systems) ( 494 , 495 ). FOAM produces surface temperatures substantially lower than those of other models and ice was accordingly thicker and slower. This is because FOAM is essentially a cloud-free model under Snowball conditions, accounting for its anomalous resistance to deglaciation at a geologically feasible CO 2 level ( 80 , 82 , 228 ).

The Snowball atmosphere has been simulated in six different GCMs under identically prescribed, 100% ice-covered conditions ( Figs. 8 and 9 ; see figure captions for model identities) ( 80 , 81 ). To isolate differences in atmospheric behavior among the models, surface albedo was set to 0.6 everywhere, eliminating differences between ablative and snow-covered ice. The prescribed albedo is appropriate for cold, ablative, meteoric ice that is free of dust and snow ( 224 , 225 ). Topography and aerosols were set to zero, as were all greenhouse gases other than CO 2 and H 2 O. The solar constant was set to 94% (present-day value; 1285 W m −2 ), obliquity was set to 23.5°, and eccentricity was set to zero. The models were run at CO 2 equal to 0.1 and 100 mbar to simulate conditions soon after a Snowball onset and just before its termination, respectively.

Whether dynamic ice sheets are compatible with a frozen ocean was among the first issues that arose when the Snowball concept was applied to Cryogenian glacial deposits ( 217 , 221 ). The issue boils down to the following. Can the Snowball atmosphere drive a hydrologic cycle capable of building ice sheets thick enough to reach the pressure-melting point at the base of the ice sheet and thereby accommodate flow by basal sliding? The ice need not slide everywhere: Large ice sheets that accumulate slowly may largely drain through narrow ice streams ( 222 , 223 ). The time frame for ice-sheet buildup is generous ( Fig. 2A ), with the Marinoan cryochron providing a somewhat tighter bound (3.0 to 15.2 My) than its ponderous predecessor. The geological evidence for dynamic ice is equally compelling in both cryochrons.

( A ) Marinoan moraine (Smalfjord Formation) resting on a quartzite bedrock pavement bearing two sets of glacial striations (arrows) in a shallow-marine paleoenvironment at Bigganjar’ga, Varanger Peninsula, Finnmark, North Norway ( 93 , 98 , 206 ). View looking eastward; 33-cm-long hammer (circled) for scale. ( B ) Polygonal sand wedges indicating subaerial exposure on the upper surface of a Sturtian glacial tillite (Port Askaig Formation), formed when glacial ice advanced across and later retreated from the paleosouthern, subtropical marine shelf of Laurentia, Garvellach Islands, Firth of Lorn, west of Scotland ( 92 , 485 ). A. M. (Tony) Spencer is seen in the lower left. ( C ) Marinoan glacial and glaciolacustrine sequence (Wilsonbreen Formation) at Ditlovtoppen, Ny Friesland, Svalbard. Glaciolacustrine carbonates (W2) yield mass-dependent and mass-independent sulfate-oxygen isotopic evidence for evaporation of liquid water and extreme atmospheric CO 2 concentration, respectively, indicating ice-free conditions shortly before the Marinoan glacial termination ( 86 , 88 , 100 ). Glacial readvance is recorded by diamictite (W3), followed by syndeglacial cap dolostone (D1) and organic-rich shale (D2) associated with the postglacial marine inundation (Dracoisen Formation) ( 486 – 488 ). ( D ) Stratified marine periglacial carbonate diamictite from the Marinoan ice grounding-zone wedge (Ghaub Formation) on the foreslope of the Otavi Group carbonate platform (Congo craton), northern Namibia ( 96 , 97 ). The pen is 15 cm long. Parallel-laminated lutite (orange) with ice-rafted debris accumulated slowly as fallout from meltwater suspension plumes, whereas graded arenite beds (gray) lacking ice-rafted debris were deposited rapidly from turbidity currents. Carbonate detritus was generated by glacial flow across a carbonate platform developed during the Cryogenian nonglacial interlude ( Fig. 2A ). ( E ) Beds of hematite-jasper (Fe 2 O 3 + SiO 2 ) iron formation (“jaspilite”) within stratified, ice-proximal, glaciomarine diamictite of the Sayunei Formation (Rapitan Group), Iron Creek, Mackenzie Mountains, Yukon, Canada. Stratigraphic context, Fe isotopes, and cerium anomaly data imply that the iron formation accumulated in a silled and likely ice-covered basin where ferruginous deep water mixed into oxygenated meltwater sourced at an advancing ice grounding line ( 160 , 220 , 350 , 400 , 412 ). Sedimentary iron formation is widely distributed in Sturtian but not Marinoan glaciomarine sequences ( Figs. 4 and 5 ). ( F ) Seafloor barite (BaSO 4 ) precipitated at a regionally extensive horizon at the top of the Marinoan syndeglacial cap dolostone (Ravensthroat Formation), Shale Lake, Mackenzie Mountains, Northwest Territories, Canada. The coin is 2 cm in diamater. Seafloor and authigenic barite is widespread in Marinoan but not Sturtian cap carbonates ( Figs. 4 and 5 ). Triple O isotope and multiple S isotopes from Ravensthroat barite constrain atmospheric CO 2 , the size of the seawater sulfate reservoir, the elimination of the atmospheric Δ 17 O anomaly, and the time scale for cap dolostone sedimentation ( 91 ).

A variety of geological features—erratic blocks ( 205 ), striated pavements ( Fig. 7A ) ( 98 , 206 – 208 ), moraines ( 98 , 206 ), faceted and striated clasts ( 209 ), preferred clast orientation ( 92 , 102 ), periglacial loessite ( 210 ), U-shaped paleovalleys ( 95 , 211 – 213 ), glacitectonic deformation ( 102 , 214 ), and glacial seismites ( 215 )—demonstrate that Cryogenian glaciers and ice sheets were dynamic. Erosion and transport may have been achieved locally by marine ice, by means of the “sea-ice escalator” ( 109 , 216 – 218 ). However, provenance indicators and depositional facies relations imply that the dominant agents of erosion and transport were grounded ice masses and their floating extensions ( 92 , 93 , 95 – 100 , 219 , 220 ).

( A ) Summer seasonal cycle of a cryoconite hole in the sublimation zone on Canada Glacier ( Fig. 13B ), Taylor Valley, Antarctica ( 279 ). In early summer, cryoconite melts to an equilibrium depth, after which it maintains constant depth relative to the sublimating surface. Air is evolved in the hole from melting of glacial ice containing bubbles of air and from photosynthetic O 2 production. Meltwater refreezes in winter and is covered by an ice cap in summer. Cryoconite is suffused with filamentous cyanobacteria and extracellular mucilaginous polysaccharides, and the holes are also inhabited by green algae, fungi, protists, and certain bilaterian animals—nematodes, rotifers, and tardigrades ( 282 ). ( B ) Postulated diurnal cycle (0000 hour) of a cryoconite hole on the low-latitude sublimation zone of a sea glacier on Snowball Earth. Relatively high CO 2 allows the nocturnal ice cap to melt away in midafternoon. Cryoconite holes and ponds provide supraglacial habitats for Cryogenian cyanobacteria and eukaryotic algae and heterotrophs once the surface becomes sufficiently warm to retain dust exposed by sublimation ( 293 – 296 ).

( A ) Cold sublimating ice surface near Mount Howe nunatak in the Transantarctic Mountains, Antarctica, at 87°22′S latitude and 2350 m above sea level. Surface dust is removed by winds, leaving a lag of stones eroded from the nunatak. Broadband albedo, α = 0.63 ( 224 ). ( B ) Warm sublimating ice surface with cryoconite holes on Canada Glacier, a piedmont glacier in the lower Taylor Valley, MDV area, Antarctica, at 77°37′S latitude and 145 m above sea level. Warmer surface temperatures due to katabatic winds allow dust (cryoconite) to accumulate on the surface, forming dark clumps suffused with organic matter that sink to an equilibrium depth, creating holes containing meltwater in summer capped by clear bubble-free ice (see Fig. 14A ). Mucilaginous and heavily pigmented organic matter is secreted extracellularly by cold-tolerant cyanobacteria inhabiting the holes, which also support eukaryotic phototrophs and heterotrophs, including metazoans. ( C ) Capsized iceberg exposing marine ice, formed by freezing seawater at a depth exceeding ~400 m, where the ice does not incorporate bubbles because of increased solubility of air in water. Nor does the ice contain brine inclusions, and light is scattered mainly by a lattice of cracks. Consequently, spectral albedo is low, α = 0.27 ( 224 ). This is the type of ice that may have been exposed in the sublimation zone of Cryogenian sea glaciers ( Figs. 15D and 18B ).

( A ) Alpine glacier and ( B ) ice sheet with terrestrial and marine ice margins. Arrows are ice flow lines. Ablation zones with transient cryoconite holes are indicated in red. ELA, equilibrium line altitude. ( C ) Sea glacier on a Snowball aquaplanet with sublimation zone (in red) where cryoconite collects. Steady-state dynamics in a 2D ice-flow model forced by the ocean-atmosphere GCM FOAM, run under relatively warm Snowball conditions ( 167 , 216 ). Sublimation of meteoric ice (compressed snow) and melting of marine ice (frozen seawater) at low latitudes are balanced by accumulation and freeze-on, respectively, outside the inner tropics. Flow velocities are highest (compare with Fig. 8D ) in the outer tropics, and ice thickness at the equator is <80 m thinner than at the poles (compare with Fig. 16 ). Overall ice thickness is determined by the geothermal heat flux, global mean surface temperature, and thermal diffusivity of ice. Salinity and, therefore, freezing temperature of seawater depend on global ice volume. ELL, equilibrium line latitude. If volcanoes and continents were included, cryoconite would accumulate in the trans-equatorial sublimation zone (red line).

Any dust or volcanic ash that accumulates on a glacier will be advected to the ablation zone and may accumulate as “cryoconite” (dark ice-dust) on the surface of sublimative ice ( Fig. 12 ). Clumps of cryoconite absorb solar radiation and sink to an equilibrium depth of ~0.5 m ( 273 – 275 ), forming “cryoconite holes” ( Fig. 13B ). Strongly pigmented organic matter produced extracellularly by indigenous cyanobacteria constitutes ~10 weight % (wt %) of typical cryoconite ( 276 ). This organic matter contributes to the dark color of cryoconite ( 276 ) and to its cohesion if desiccated ( 277 ). Although cryoconite holes on modern polar glaciers ( Fig. 14A ) contain meltwater only in summer ( 278 , 279 ), they support ecosystems that include not only cyanobacteria ( 280 , 281 ) but also eukaryotic green algae, fungi, ciliates, and certain metazoans, typically nematodes, rotifers, and tardigrades ( 282 , 283 ). We will return to the subject of cryoconite and its role in the dynamics and ecology of oceanic ice on Snowball Earth.

Dry valleys, warmed by katabatic winds from adjacent ice sheets, were likely refugia for cold- and desiccation-tolerant microbial phototrophs and heterotrophs on Snowball Earth ( 261 , 262 ). In 1971, cyanobacteria were found inhabiting rocky crevices at 2000 m elevation in the Transantarctic Mountains, <360 km from the South Pole ( 263 ). In the MDV, 1000 km to the north and 1.8 km lower in elevation, cyanobacteria inhabit rocky soils ( 264 ), lake ice ( 265 ), and ice-covered hypersaline lakes ( 266 , 267 ) in areas that rarely touch the melting point ( 268 ). Diatoms and ciliates constitute a small eukaryotic component of plankton in MDV brine lakes ( 269 ), and benthic diatom mats form stromatolites in MDV meltwater streams ( 270 ). Similar ecosystems are found in the High Arctic ( 271 , 272 ). However, diatoms do not appear in the fossil record until the Mesozoic era.

An analogous situation is inferred from totally different evidence in the Marinoan of East Svalbard ( Fig. 5A ). Periglacial lacustrine limestone ( 100 – 102 ), associated with dolostone and sandwiched between massive glacigenic diamictites ( Fig. 7C ), contains trace sulfate bearing the largest mass-independent oxygen isotope anomaly (Δ 17 O ≥ −1.6‰) in the terrestrial record ( 86 ). The magnitude of the anomaly can be accounted for by high atmospheric CO 2 level combined with low rates of photosynthesis-respiration ( 86 , 88 , 163 ). Moreover, there is a strong mass-dependent isotopic enrichment of oxygen in the host carbonate, δ 18 O ≤ +15‰ ( 86 ), most likely the result of evaporation. Sublimation of a permanently ice-covered lake should drive sub-ice water isotopically lighter, not heavier, because of ice-water equilibrium fractionation (δ 18 O = 1.003) ( 259 ) and quantitative sublimation of ice. Accordingly, either the Marinoan lake was not permanently ice-covered or the meltwater that fed it underwent strong evaporation as it flowed into the lake basin from adjacent glaciers. In contrast, modern hypersaline lake waters in the Antarctic McMurdo Dry Valleys (MDV) are isotopically light, close to the compositions of the bordering glaciers ( 260 ). This suggests that the opposed isotopic effects of evaporation and sublimation are nearly balanced in the MDV, whereas the late Marinoan cryochron in East Svalbard was more strongly evaporative.

Geological evidence supports the existence of ice-free land areas on Snowball Earth with surface temperatures near the melting point, as predicted by the LMDz-GRISLI climate model ( 100 ) when CO 2 is high ( Fig. 10D ). During the Marinoan cryochron in South Australia ( Fig. 5A ), a low-latitude periglacial block field was invaded by a sand sea, apparently under the influence of katabatic (paleo-north-northwesterly) winds ( 258 ). The block field and basal sandstone beds host deep (≤3 m) sand wedges, indicating subaerial exposure in a periglacial environment with strong seasonality ( 230 – 232 ). Ductile deformation of beds associated with sand wedges and the presence of surface meltwater channels indicate that the annual mean surface temperature should have been ≥268 K ( 232 ) at the known paleomagnetic latitude of 7° to 14° ( 33 ). The sand sea apparently formed toward the end of the Marinoan cryochron, whereas a coeval ice sheet and associated katabatic winds are inferred to account for south-southeastward–directed dune migration ( 258 ), in the zone of easterly trade winds.

Same coupled atmosphere–ice-sheet model as in Fig. 10 , with CO 2 = 20 mbar. Model equilibrated to northern-hemisphere ( A ) warm (WSO-N) and ( B ) cold (CSO-N) summer orbits. Orbits were switched every 10 ky. Scale bars are ice thickness (in meters), and brown areas are ice-free. ( C ) Difference in local ice-sheet mass balance, with red and blue indicating positive and negative mass balance, respectively, in the WSO-N relative to CSO-N. Note the hemispheric asymmetry as expected for precession, but areas of positive and negative mass balance coexist in both hemispheres at low latitudes, expressing the sensitive response of the hydrologic cycle. Positive mass balance in the war-summer hemisphere is related to prescribed change in eccentricity. ( D ) Expanded sector (magenta box) under CSO-N with white lines indicating ice margins under WSO-N for comparison. Ice margins migrate <5° (550 km) on the precessional time scale.

To investigate the sensitivity of ice sheets on Snowball Earth to orbital forcing, integrations at each CO 2 level have been performed in which warm and cold summer orbits are switched reciprocally in each hemisphere every 10 ky ( 100 ). Precession-like forcing is chosen to investigate ice-sheet response in the tropics, where obliquity forcing is found to be weakest, as expected. Switching between orbital extremes, while not realistic, is computationally efficient and demonstrates that the ice-sheet response can outpace sinusoidal precession of the ecliptic. The atmosphere and ice-sheet models are coupled every 10 ky, which may dampen the ice-sheet response slightly compared with more frequent coupling, because of the delayed ice-elevation feedback. The observed response is strongly CO 2 -dependent. At the lowest level (0.1 mbar), the ice sheets are unresponsive to orbital switching on a precessional time scale. At intermediate levels, the ice response is strong, and tropical ice-sheet margins migrate as far as 5° (550 km) latitude at 20 mbar ( Fig. 11D ). Ice volume increases in the cold-summer hemisphere, as expected for precessional forcing, but the response is spatially complex with simultaneous positive and negative mass balance in different areas of each hemisphere ( Fig. 11C ). At 100 mbar, the response weakens again because the amount of remaining ice is small ( Fig. 10D ). The sensitivity of ice-sheet margins to precession-like forcing at intermediate CO 2 supports a possible orbital origin for depositional cycles in Cryogenian glacial-periglacial sequences ( 100 ). It is also consistent with a U-Pb zircon age of 640.3 ± 0.4 Ma, 4 to 6 My older than the Marinoan termination, from a tuff at a stratigraphically intermediate level within the ice grounding-zone wedge on the Congo foreslope ( 83 ). The magnitude of the flooding events associated with both Cryogenian glacial terminations implies the existence of more ice than modeled at 100 mbar of CO 2 ( Fig. 10D ). Additional modeling is needed to see whether deglaciation can be triggered at lower CO 2 and a larger ice-sheet volume, possibly by lowering the equatorial oceanic ice albedo, as discussed in the “Cryoconite holes and ponds” section.

Results from experiments with the atmospheric component of GCM LMDz, coupled to ice-sheet model GRISLI, in Cryogenian paleogeography with prescribed orography under Snowball conditions with present-day orbit and CO 2 = 0.1 mbar ( A ), 20 mbar ( B ), 50 mbar ( C ), and 100 mbar ( D ). Scale bar is ice-sheet thickness (in kilometers), and brown areas are ice-free. Red circles indicate the study location in Svalbard. Large base-level rise associated with Marinoan deglaciation may imply that more ice was available for melting than in (D), consistent with terminal deglaciation at less than 100 mbar of CO 2 .

To investigate the ice-sheet response to orbital forcing and CO 2 variation on Snowball Earth, experiments have been conducted with a 3D atmospheric model coupled to an ice-sheet model in Marinoan paleogeography ( 31 ) with prescribed mountain ranges ( 100 ). The oceanic component of the climate model is turned off because the ocean is prescriptively ice-covered. The atmospheric model (LMDz) ( 256 ) and ice-sheet model [Grenoble Ice Shelf and Land Ice (GRISLI)] ( 257 ) are the same ones prescribed in an earlier study ( 165 ) with 0.33 mbar of CO 2 , 94% of present solar luminosity, and the present orbital configuration and day length. In response, the model oceans are frozen, and ice sheets that had been “seeded” on the mountains extend over all tropical and most mid-latitude continents within 0.2 My, leaving bare only some coastal strips ( 165 ). The new experiments ( 100 ) seek to compare equilibrium ice-sheet volumes at different CO 2 levels, as well as the response of the ice sheets to orbital-like forcing at each level. To expedite ice-sheet initialization, ablation is eliminated for the first 500 ky, prescribed with present orbital parameters and low CO 2 (0.1 mbar). The models are incrementally coupled every 10 ky. Ablation is introduced at 500 ky, after which the ice sheets wane toward equilibrium volumes at prescribed CO 2 levels of 0.1, 20, 50, and 100 mbar. The coupling interval for this phase is 100 ky. At 0.1 mbar, the model converges on an ice volume close to that found in an earlier study ( 165 ) within 600 ky after ablation is introduced. The presence of mountains causes ice to overrun the low-latitude zone of net ablation. At each progressively higher CO 2 level, the equilibrium ice volume diminishes ( Fig. 10 ). At 100 mbar, it is <20% of the equilibrium ice volume at 0.1 mbar and is largely controlled by topography ( 100 ).

Likewise, during the Marinoan cryochron, grounded ice flowed off the subtropical northeastern (present southwestern) margin of the Congo craton ( Fig. 5A ), leaving a compound ice grounding-zone wedge, the Ghaub Formation, on the steep foreslope of a wide carbonate shelf ( 96 , 97 ). The wedge formed where inland ice flowed across a grounding line into a floating ice shelf. The wedge is composed of interfingered massive and stratified carbonate diamictites, with a terminal ferruginous drape (Bethanis Member) crowded with ice-rafted debris. The massive diamictites include melt-out and rain-out deposits, locally channelized by well-sorted deposits or deformed by overriding ice ( 96 ). The stratified diamictites ( Fig. 7D ) accumulated subaqueously in a marine setting and are products of three simultaneous depositional processes: fall-out from fine-grained suspension plumes, mass flows (turbidites and debrites), and ice-rafted dropstones. Ice rafting could have been accomplished by icebergs or by a continuous ice shelf, except for the terminal drape, in which nested dropstones are more consistent with iceberg rafting ( 250 ). Well-sorted, carbonate sand and gravel form channels within massive diamictites and fans in stratified diamictites. The channels and fans contain dropstones and attest to subglacial meltwater flow and discharge at the grounding line, respectively. The cycles are typically asymmetric, recording progressive grounding-line advances punctuated by abrupt retreats. The grounding-zone wedge rests on an apparent ice-cut surface, implying an older ice maximum ( 222 ). The wedge pinches out upslope and tapers downslope but continues for hundreds of kilometers along strike. The grounding-line migrations were apparently limited to the width of the wedge, consistent with the steep inclination of the subglacial surface and the absence of reverse bed slopes in the foreslope area ( 97 , 251 ).

Cryogenian glacial and glacial marine deposits commonly display vertical alternations of distal and more proximal deposits that are interpreted to reflect repetitive advances and retreats of ice margins or grounding lines. The scale of these apparent cycles is broadly comparable to that of orbital cycles in Cenozoic periglacial sequences, but their frequency in the Cryogenian is unknown. During the Sturtian cryochron on the southern subtropical margin of Laurentia ( Fig. 5B ), for example, grounded ice flowing from the continental interior advanced across a former carbonate-rich, shallow-water marine shelf. The ice left stacked tabular bodies of ice-proximal (rain-out) and ice-contact (melt-out) diamictites, the Port Askaig Formation, with deformed zones attributable to overriding ice ( 92 ). No fewer than 13 of the diamictite bodies have their upper surfaces ornamented by polygonal sand wedges ( Fig. 7C ). The sand wedges indicate subaerial exposure in a periglacial environment, evincing withdrawal of the ice that deposited the diamictite on which the sand wedges developed ( 92 ). In the absence of a detailed chronology, we cannot say whether the migrations of the ice margin were externally forced (for example, orbital) or an expression of ice-sheet dynamics. Polythermal ice sheets in cold climates may undergo repeated surge-retreat cycles due to internal dynamics without external forcing ( 248 , 249 ).

OCEANIC ICE AND THE EVOLUTION OF CRYOGENIAN MARINE LIFE

Snowball Earth is essentially an oceanographic phenomenon. Its onset is defined when the tropical ocean freezes over, and its termination is defined when the equatorial ice shelf finally divides and collapses. Small areas of open ocean are unsustainable because the sea ice becomes hundreds of meters thick within a few thousand years, due to the albedo-driven cold surface temperatures. Consequently, the ice spreads gravitationally and fills in any area that is not physically restricted (116, 167, 168, 216, 224). The term “sea glacier” (216, 224) describes this floating ice mass (Figs. 8, C and D, and 12C), which flows toward the equatorial zone of net ablation from higher latitudes of net accumulation (Fig. 9, E and F). Likewise, it is difficult to maintain areas of oceanic ice sufficiently thin for sub-ice phototrophy, <20 m of clear ice, particularly in the coldest early part of a cryochron (224, 225, 284–288). The most favorable conditions for thin oceanic ice exist in hydrothermal areas (289) and marine embayments into low-albedo ice-free land areas (290–292).

Thin oceanic ice (or open water) is not a prerequisite for phototrophy if meltwater existed on the ice surface. Impressed by the existence of extensive supraglacial cryoconite ponds and associated microbial mats on the McMurdo (78°S) and Ward Hunt (83°N) ice shelves, Vincent and colleagues (293–295) postulated that similar ecosystems on Snowball Earth “would have provided refugia for the survival, growth and evolution of a variety of organisms, including multicellular eukaryotes.” The proposal was reinforced when modeling (100) indicated that ice-free land areas on Snowball Earth widen from the paleoequator as CO 2 rises (Fig. 10). Moreover, those source areas of terrigenous dust are situated in the same zone where surface winds associated with the Hadley cells are strongest (Fig. 9, A to D). The Snowball troposphere was dusty, and dust trapped anywhere on the sea glacier or on ice sheets feeding the sea glacier would be carried by glacial flow to the sublimative surface of meteoric ice in the equatorial zone (Fig. 12C). In Cryogenian paleogeography, the sea glacier sublimation zone is ~6 × 107 km2 or about 12% of global surface area (296). Whereas cryoconite holes and ponds in the polar regions freeze solid in winter (Fig. 14A), those in the equatorial zone of Snowball Earth may have always contained meltwater (Fig. 14B), except during the earliest stages of a cryochron when sublimative surfaces may have been too cold for dust retention (Fig. 13A).

We begin the assessment of supraglacial refugia hypothesis by Vincent and co-workers (293–295) by briefly reviewing what is known from molecular and body fossil evidence about pre-Sturtian and pre-Marinoan marine life, with emphasis on purported crown groups, which, by definition, are lineages that survived the cryochron(s), and all subsequent vicissitudes. We then review attempts to find geologically acceptable climate-model states in which the tropical or equatorial ocean remains ice-free. Next, we sketch sea-glacier dynamics and its response to surface warming, based on 2D and 3D models. Finally, we consider the timing and extent of cryoconite accumulation and its potential climatic, geochemical, and evolutionary consequences.

Pre-Sturtian and Cryogenian fossil record Cellular fossils and molecular phylogenetics indicate that cyanobacteria, including those with cellular differentiation, evolved more than 109 years before the Cryogenian (297–300). Low ratios of eukaryotic-to-bacterial biomarkers from indigenous bitumens and oils imply that bacteria were the dominant primary producers in pre-Sturtian oceans (49, 301). As for eukaryotic primary producers, red algae and possibly green algae, including multicellular forms, are known from the pre-Sturtian cellular fossil record (298, 302–305). Molecular (sterane) biomarkers suggest that green algae supplanted red algae as the dominant eukaryotic phototrophs sometime between the late Tonian and late Cryogenian (48, 49, 301). Among eukaryotic heterotrophs, vase-shaped microfossils (VSMs) resembling extant amoebozoans and rhizarians are widely preserved in pre-Sturtian strata around 740 Ma (306–309), and various protistan morphotypes including VSMs are found in nonglacial strata between the cryochrons (310–313). Molecular clocks predict that stem-group metazoans predated the Sturtian cryochron (47), and sterane biomarkers suggest that a metazoan crown group, demosponges, evolved before the Marinoan cryochron [(46–48); but see the study by Brocks and Butterfield (314)]. A Sturtian origin for crown-group metazoans is estimated by a molecular phylogenetic “clock” (47, 315), although weak pre-Cambrian calibration compromises the accuracy of this estimate (108). The fossil record in total is too coarse to correlate extinctions or originations with cryochrons, but the Cryogenian stands out as an anomalous period of low total and within-assemblage eukaryotic diversity (103). After the Cryogenian ended, acritarch diversity increased sharply (316, 317), as perhaps did that of benthic macroalgae (106, 107, 318). The fossil record and molecular phylogeny together indicate that multiple clades of eukaryotic algae and heterotrophs, both single-celled and multicellular, not only survived the Cryogenian glaciations but may have significantly evolved during that period (319–321).

Waterbelt solutions There have been concerted efforts to find climate-model solutions that satisfy basic inferences from Cryogenian geology—dynamic ice sheets that reach sea level in the paleotropics (Fig. 5)—while maintaining a finite zone of open water in the warmest area. Some of these efforts were motivated by a perception that the “hard-snowball” hypothesis is implausible in light of the fossil record (111, 112, 114, 322, 323). “Waterbelt” (116) is a general term for these “loophole” (322) solutions, which have been less accurately called “slushball” or “soft-snowball” solutions in the literature. The modeling task is a difficult one. First, the solutions, by their nature, lie close to the Snowball bifurcation (Fig. 1), yet they must resist falling irrevocably into the Snowball state for millions to tens of millions of years (Fig. 2). This is a tall order, given stochastic, orbital, tectonic, and paleogeographic forcings (116). Second, a large hysteresis must exist between the Waterbelt and nonglacial states (116) to satisfy the geologically observed abrupt deglaciations and attendant geochemical anomalies—cap carbonates (27, 29, 84, 154, 324, 325), proxy indicators of high CO 2 (84–88, 90), and spikes in weathering (60, 89, 326). Third, as applied to the Sturtian cryochron, solutions must be compatible with deep-ocean ferruginous anoxia, given widespread synglacial nonvolcanic iron formations (Figs. 5B and 7E). This is a challenge because a narrow tropical ocean will experience intense wind-driven ventilation (236, 237), and the remote ice-covered regions will lack organic productivity, reducing the demand for oxygen by aerobic respiration in those areas. However, the model solutions are interesting in their own right, independent of the needs of Cryogenian geology. The most-cited Waterbelt solution, HCBP00 (327), emerged from a 2D energy-balance model coupled to a dynamic ice-sheet model, with a paleogeography in which a high-latitude supercontinent has promontories and large islands that extend across the deep tropics. Long integration times allow orbital forcing to be included. By incrementally lowering the CO 2 radiative forcing, a Snowball bifurcation is found at which ice sheets abruptly extend to all latitudes. The fully glaciated response from the coupled model was then prescribed in an atmospheric GCM (Genesis 2) with a mixed-layer ocean (that is, no ocean dynamics) and nondynamic sea ice. Over a limited range of CO 2 and continental freeboard, the tropical ice sheets coexist stably with sea-ice edges at ~25° latitude (327). The solution was criticized for lacking sea-ice dynamics (328, 329). Sea-ice dynamics facilitate ice-line advance in the mid-latitudes where the Coriolis effect drives sea ice equatorward under the influence of westerly winds. The concern was left unresolved because the wind field in the dynamic sea-ice model (329) was imported from a GCM (FOAM) response to Cryogenian paleogeography in the absence of ice. A low-latitude ice margin would produce a much stronger wind field (236, 237), in which Coriolis forcing under the influence of the trade winds might actually retard sea-ice advance. Another concern with the HCBP00 solution is that the mid- to high-latitude sea-ice caps are arbitrarily limited to 10 m of maximum thickness, and therefore, gravitational flow (116, 167, 168, 216) is excluded. Moreover, an artificial heat source was introduced to limit ice thickness (327), and this heat source arbitrarily retards ice-cap growth. Additional simulations showed that the sea-ice margins in the HCBP00 solution retreat poleward in response to even modest increases in CO 2 (330), simulating the loss of weathering by the ice-covered continents. The hysteresis demanded by the records of Cryogenian deglaciation is not present. It was subsequently found that low-latitude ice sheets do not develop when the tropical ocean is ice-free in a coupled ocean-atmosphere GCM (ECHAM5/MPI-OM), even when mountain topography is prescribed (166). A different class of Waterbelt solutions, called “Jormungand” (236), was revealed by simulations with atmospheric GCMs (CAM and ECHAM5) and coupled atmosphere-ocean GCMs, with (CCSM3 and CCSM4) and without (ECHAM5/MPI-OM) sea-ice dynamics. It was found that by prescribing a large difference in broadband albedo between ablative (0.55) and snow-covered (0.79) ice, sea-ice margins are stable at 5° to 15° latitude (236). As the floating ice margins enter the tropics in response to weakened radiative forcing, the ablation zones widen as they encroach upon the subsiding limbs of the intensified Hadley cells. This lowers the zonal and planetary albedos, stabilizing the ice margins. The narrow seaway migrates nearly its own width back and forth across the equator with the seasons. Ice sheets develop on elevated continents in the equatorial zone where, unlike Snowball Earth, there is a large excess of precipitation over evaporation (166, 236). Strong hysteresis between Jormungand and nonglacial states has been found in aquaplanet atmosphere-only GCMs, although less so than for the Snowball state (236), but more work is needed to clarify the impact of ocean dynamics and continents on Jormungand hysteresis. The low-latitude sea-ice margins destabilize to Snowball states when sea-ice dynamics are switched on in ECHAM5-MPI-OM (20) but not in CCSM3 and CCSM4, which include sea-ice dynamics (19, 332). Orbital forcing and sea-glacier flow have yet to be investigated in the Jormungand state. A third Waterbelt solution, BR15 (237), was uncovered using a coupled atmosphere-ocean-sea ice GCM (MITgcm) with simplified paleogeographies. As CO 2 is incrementally lowered, the sea-ice margins stabilize at 21° to 30° because of wind-driven ocean heat transport, which intensifies as the margins converge (consistent with Jormungand), creating a negative feedback (237). As with Jormungand, equatorial ice sheets are compatible with the BR15 solution (166). A potential destabilizing process involves boundary-layer temperature inversions in the winter hemisphere, where the surface becomes very cold through radiation (81, 226–228, 331). Inversions decouple the ocean from winds, disabling the wind-driven negative feedback. Atmospheric GCMs with high vertical resolution and realistic paleogeography are needed to resolve this issue.

Grounding-line crack systems Sea-glacier flow appears to foreclose the possibility of widespread photosynthesis beneath tropical sea ice, but it suggests another setting where liquid water and sunlight might intercept. Shear cracks would perpetually develop where a fast-flowing sea glacier is in contact with landfast ice (Fig. 17). Where inland ice sheets drained into the ocean through outlet ice streams, they might simply merge with the sea glacier without deep cracks. But where the inland ice was frozen to the bed, away from ice streams, shear stress would be relieved by deep crack systems in which water pressure would hold opposing ice cliffs apart. A modern analog occurs on the north side of the Pine Island Ice Shelf (West Antarctica) near the calving front (Fig. 17A). The ice shelf is ~0.5 km thick at this point (336), and a dextral shear couple exists between landfast ice and fast-flowing (~2.8 km year−1) shelf ice. The cracks are more open close to the calving front (Fig. 17A), which presumably would not exist on a Snowball Earth until the terminal deglaciation. Fig. 17 Shear cracks on ice shelves. (A) Dextral shear produces a crack system where fast-flowing (~2.8 km year−1) 0.5-km-thick shelf ice abuts grounded ice on the north side of the Pine Island Ice Shelf, West Antarctica. Crack system is best developed within 20 km of the calving front. Arrows indicate ice-shelf flow direction, and tacks indicate landfast ice. Satellite imagery courtesy of NASA/GSFC/METI/ERSDAC/JAROS U.S./Japan ASTER team. (B) Thickness of a sea glacier on Snowball Earth (Figs. 8C and 16) implies that cracks were deeply recessed, weakly illuminated, and more important as conduits for air-sea gas exchange than for phototrophy. The seawater that filled a newly formed crack would encounter the cold atmosphere and freeze at the surface, forming new sea ice with inclusions of seawater that would salinate by progressive freezing as the sea ice thickened and chilled. Brine inclusions in modern sea ice provide habitats for prokaryotic and eukaryotic phototrophs (mainly diatoms), as well as heterotrophic protistans and metazoans (337). The organisms must be tolerant not only of hypersalinity but also of hyperoxia, because photosynthetically produced O 2 cannot diffuse out of a closed brine inclusion. The thickness of sea-glacier ice (Figs. 8C and 16, A and B) is such that sea ice in cracks would reside in 0.1-km-deep “canyons” and receive direct sunlight for only a short period of the day (Fig. 17B), if at all, unless they were fortuitously oriented in the ecliptic plane.

Dust sources and accumulation rates There were three sources of dust on Snowball Earth, volcanic, detrital, and cosmic (78). The modern production rate of volcanic tephra (ash and coarser-grained ejecta) amounts to an accumulation rate, if spread evenly over the globe, of roughly 10−6 m year−1 (78, 338). The terrestrial volcanic flux would be somewhat reduced on Snowball Earth due to loading by ice sheets (339), but this effect would wane over time as ice sheets contracted with CO 2 rise (Fig. 10). The modern global average accumulation rate of detrital dust is roughly 5 × 10−7 m year−1 and was 2 to 20 times higher (10−6 to 10−5 m year−1) at the Last Glacial Maximum (LGM), mainly due to decreased vegetation (78, 340). The LGM dust flux is a reasonable minimum estimate for Snowball Earth (77–81), given extensive ice-free land area (24, 92, 100, 230, 232, 258, 341, 342), arid and poorly vegetated soils produced by intense cryogenic weathering associated with large diurnal and seasonal temperature fluctuations (77, 343), production of loess through the grinding action of glaciers charged with rock debris (344), and strong summer and katabatic winds (81, 226–228, 258, 342). In comparison, the modern global average accumulation rate of cosmic dust, 1.5 × 10−10 m year−1, is negligible (78). There is no reason to suspect that the Cryogenian cosmic dust flux was significantly higher than modern values (345). Taking the LGM flux of detrital dust alone amounts to a global average accumulation rate of 1 to 10 m My−1. Considering the duration of a Snowball Earth, where did so much dust actually accumulate?

Cryoconite holes and ponds Dust that is trapped in the accumulation zones of ice sheets and the sea glacier will be buried by new snow and entrained in meteoric ice (Fig. 12). Roughly half the ice that accumulates in ice sheets drains into the sea glacier, whereas the rest deposits its dust load in peripheral moraines (Fig. 10). The total ice-sheet area shrinks from 0.8 to 0.2 of continental (including shelf) area over a cryochron (100) or from 0.32 to 0.08 of global surface area. Assuming that the sea glacier occupies 0.6 of global surface area, the total area of accumulation that will be advected to the sublimation zone of the sea glacier will be 0.76 to 0.64 of global surface area. Because the sea-glacier sublimation area is 0.12 of global surface area in Cryogenian paleogeography, the dust accumulation rate in the sublimation zone, assuming zero recycling by winds (or meltwater flushing, as discussed in the next section), will be 6.3 to 5.3 times the global average dust flux. Taking the conservative LGM dust flux of 1 to 10 m My−1 leads to average accumulation rates in the sea-glacier sublimation zone of roughly 6 to 60 m My−1. The total thickness of dust that would have accumulated over the 58-My Sturtian cryochron is 0.35 to 3.5 km. It seems improbable that the equatorial sea glacier could mechanically support a supraglacial moraine of this thickness. Hence, let us examine the scenario more closely. At the onset of a Snowball Earth, when sublimative surfaces are extremely cold, dust will not stick but will be lofted by winds and filtered out of the air by the firn (uncompacted snow) in accumulation zones (346). Modern examples exist in the “blue ice” areas of the Transantarctic Mountains (Fig. 13A), where both dust and rock fragments (including meteorites) are advected to the sublimation surface, but only stones too large for aeolian transport accumulate. Cryoconite and cryoconite holes are absent on such a cold surface. In the early phase of a Snowball cryochron, cryoconite holes may only develop in equatorial marine embayments, where sublimative ice is flanked by low-albedo ice-free land (290–292). Elsewhere, dust will be continuously recycled, poleward by wind and equatorward by ice, while accumulating within the sea glacier, rather than on its sublimative surface. As atmospheric CO 2 rises and the subsurface ice becomes dustier, the ablative surface warms and dust begins to stick. The ice albedo drops rapidly. As surface dust accumulates, it clumps and absorbs sunlight, creating meltwater films in which cyanobacteria grow. They secrete heavily pigmented extracellular polysaccharide that darkens the dust, increases its wind resistance, and contributes to mass wasting of the ice (276, 277, 347). Clumps of dark dust sink to an equilibrium depth of 0.4 to 0.6 m in the ice, forming cryoconite holes (Fig. 13B). Whereas polar cryoconite holes contain liquid water only in summer and may be permanently ice-capped (Fig. 14A), cryoconite holes on the equatorial sea glacier may contain liquid water throughout the year. They will be insulated at night by caps of ice and exposed to the atmosphere when afternoon temperature at equinox reaches the melting point (Fig. 14B). The absence of winter freeze-up makes cryoconite holes on an equatorial sea glacier a less stringent habitat than those on modern polar ice shelves and glaciers. The limitation to growth is the availability of mineral nutrients from the dust. Dominance of cyanobacteria is a hallmark of oligotrophy in cold environments (348). If the rate of dust accumulation on the sublimation surface of the sea glacier is anywhere near 6 to 60 m My−1, then cryoconite will rapidly (on a Snowball time scale) saturate the surface (77, 78, 167, 168). Cryoconite holes will coalesce to form meter-deep cryoconite ponds (261, 293–295). The sublimation zone area in Cryogenian paleogeography is roughly 6 × 107 km2 or 0.12 of global surface area (Fig. 18A). Solar energy absorbed by a semicontinuous layer of cryoconite has a marked effect on ablation-zone ice thickness (Fig. 18B), particularly near the snowline latitude where the dusty surface is first exposed (168). In a 1D (meridional) ice-atmosphere-dust climate model, operating within the framework of a 1D energy-balance model, sublimation zone ice thicknesses are about 0.2 and 0.1 km for dust fluxes of 10−6 and 10−5 m year−1, respectively (168). For comparison, the equivalent ice thickness in the same model in the absence of cryoconite is 0.74 km (168). In the cryoconite-rich zone, the temperature difference between the top and bottom of the sea glacier is nearly zero; thus, the equilibrium ice thickness in the absence of flow would be zero. The modeled ice thicknesses of 0.1 to 0.2 km are maintained by the inflow of cold ice (168). The inflows are characterized by quasi-periodic, reciprocating, semicentennial, ice surges (168). Fig. 18 Cryoconite ponds and cryoconite meltwater flushing on Snowball Earth ( 297 ). (A) Global paleogeography during the Sturtian cryochron at 680 Ma (34). Brown continental areas schematically indicate dry-valley dust sources (100). Gray areas in the low-latitude sublimation zone of the sea glacier are schematic cryoconite ponds. (B) Snapshot of a 2D ice-flow model of a Snowball Earth with a global dust accumulation rate of 10 m My−1 (168). Dark cryoconite and exposed marine ice make sublimative ice warm and thin. Hemispheric asymmetry at low latitude reflects unsteady, quasi-periodic, reciprocal ice surges from the respective subtropics in the model. Flushing of meltwater and cryoconite through moulins is shown schematically. The apparent ice walls at 18° latitude have actual slopes of 1 m km−1. Organic production is nearly balanced by aerobic respiration within cryoconite holes and ponds. Flushed cryoconite organic matter is subject to anaerobic respiration in the water and sediment columns. Fe(III) and sulfate are sourced from the flushed cryoconite and sub–ice-sheet weathering. If anaerobic respiration is incomplete, organic matter is buried and O 2 is added to the atmosphere. SL, sea level.

Meltwater flushing and the cryoconite thermostat How could a Cryogenian Snowball Earth persist for 5 × 107 years encircled by an equatorial cryoconite moraine hundreds to thousands of meters thick? Such a moraine would significantly lower the planetary albedo, but it would also shield the underlying ice from penetrative radiation. One possibility is that overthickened areas of the moraine would become gravitationally unstable, leading to localized ice collapses and dumping of morainal sediment into the subglacial ocean. The collapsed areas would quickly “heal” by glacial inflow, resulting in a thinner moraine. Because the morainal load is spread out, collapses should not occur until the sea glacier has become quite thin. An alternative stabilizing feedback that does not require a thick moraine involves meltwater flushing through enlarged cracks, called moulins, and consequent cryoconite cleansing (168). As cryoconite accumulates, the rate of meltwater production rises. Drainage systems develop, linking the cryoconite ponds to flushing conduits, or moulins—subvertical shafts that originate as cracks and are maintained by the latent heat of falling meltwater and the hydrostatic pressure of seawater acting over 90% of their depth. The drainage systems cleanse the ice surface of cryoconite and flush it into the subglacial ocean (Fig. 18B). This raises the ice albedo and reduces the meltwater production rate. If the dust flux wanes, then meltwater production slows and cryoconite accumulates. If the dust flux waxes, then meltwater production quickens and cryoconite is removed. The resulting “cryoconite thermostat” (168) maintains a relatively warm and thin equatorial sea glacier (compare Figs. 12C and 18B) but is incapable of triggering terminal deglaciation at low CO 2 . A critical CO 2 threshold is still required, although it will be substantially less than if cryoconite was absent. In the sea-glacier model with cryoconite (Fig. 18B), marine ice is exposed in the sublimation zone (unlike Fig. 12C), as required by a steady-state hydrologic cycle (Fig. 15D).

Meltwater flushing, the carbon cycle, and atmospheric oxygen The organic contents of modern cryoconite on Himalayan, Tibetan, and Arctic glaciers range from 3 to 13 wt %, with the highest average contents (11 wt %) in the Arctic (276). Within cryoconite holes and ponds, organic production is nearly balanced by aerobic respiration over the seasonal cycle (278, 349). But what happens to the organic content of cryoconite that is flushed into the subglacial ocean (Fig. 18B)? That ocean is generally assumed to have contained little dissolved oxygen (24, 74, 350), consistent with the wide distribution of Sturtian synglacial iron formation (Figs. 4B and 5B). Pre-Sturtian deep water was largely ferruginous, with intermittent to persistent euxinia in highly productive coastal areas (351–354). In the Snowball ocean, O 2 influx at cracks and moulins would have been offset by consumption of O 2 by seafloor weathering and submarine volcanic outgassing (74). In the absence of O 2 , respiration depends on sulfate and Fe(III) as terminal electron acceptors. The subglacial ocean had two sources of sulfate and Fe(III). The injection of meltwater generated beneath ice sheets (Fig. 15D) delivered dissolved sulfate and suspended Fe(III) as products of oxidative subglacial weathering (171). Cryoconite flushing delivered sulfate derived from volcanic aerosols and Fe(III) from detrital dust. If the fluxes of sulfate and Fe(III) were inadequate to respire all the flushed cryoconite organic matter, in the water column or in the sediment, then organic burial would have created a source of atmospheric O 2 (Fig. 18B). A source of O 2 was needed to meet the consumptive demands of subaerial volcanic outgassing and rock weathering, because the absence of mass-independent S isotope fractionation (δ33S ≥ 0.3‰) in Cryogenian sediments (355) implies that the Snowball atmosphere did not become anoxic. It has been hypothesized that Snowball glaciations of Siderian (2.5 to 2.3 Ga) and Cryogenian age (Fig. 2B) were responsible for irreversible increases in atmospheric O 2 (356). Cryoconite flushing and resultant organic burial are processes by which this might have been achieved (296). Although there is some proxy support for stepwise increases in atmospheric O 2 coincident with Cryogenian glaciations [(357–360); but see the study by Blamey et al. (361)], the existing record of atmospheric oxygenation between 0.8 and 0.4 Ga overall remains inadequate to confidently defend or refute the hypothesis (356) that Cryogenian glaciation drove atmospheric O 2 irreversibly from a Proterozoic to a Phanerozoic steady state.

Cryoconite flushing and volcanic ash deposition Discrete layers of volcanic ash (Fig. 19A) are occasionally found within stratified glaciomarine deposits of Cryogenian age (32, 56, 58, 59, 83, 362–364). Where intrabasinal volcanism is absent, the ash layers are assumed to be far-traveled and subaerially erupted. It has been argued that the deposition of these layers indicates ice-free conditions (83). But what if the ash was deposited first on ice and then glacially advected to the low-latitude sublimation surface of the sea glacier (Fig. 12C), where it resided until flushed through a moulin into the subglacial ocean (Fig. 18B)? Glacial flow is nearly linear, so a patch or streamer of volcanic ash would maintain its integrity in transport, becoming more concentrated after it reaches the edge of the sublimation zone (168). The ash-rich cryoconite would be susceptible to dispersal by ocean currents once it is fl