Research ArticleCLIMATOLOGY

Snowball Earth climate dynamics and Cryogenian geology-geobiology

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Science Advances  08 Nov 2017:
Vol. 3, no. 11, e1600983
DOI: 10.1126/sciadv.1600983


  • Fig. 1 Generic bifurcation diagram illustrating the Snowball Earth hysteresis.

    Ice-line latitude as a function of solar or CO2 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 CO2 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 CO2 combined with low surface albedo creates a torrid greenhouse climate. Intense silicate weathering and carbon burial lower atmospheric CO2 (in 107 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 CO2 radiative forcing (c and d). The ocean is predicted to undergo severe acidification and deacidification in response to the CO2 hysteresis. Qualitatively similar hysteresis is found in 3D general circulation models (GCMs). Pco2, partial pressure of CO2; wrt, with respect to.

  • Fig. 2 Glacial epochs on Earth since 3.0 Ga.

    (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.

  • Fig. 3 Early metazoan phylogeny, the geologic time scale, and Neoproterozoic glacial epochs.

    Simplified topology of the metazoan radiation (315), based on a phylogeny and concatenated, invertebrate-calibrated, molecular clock estimates (47) and glacial chronology (Table 1). Sterane biomarkers in the South Oman Salt Basin are interpreted to be products of demosponges, appearing no later than the clement interlude between the Sturtian and Marinoan glaciations (46, 48).

  • Fig. 4 Present distribution of Cryogenian glacial-periglacial deposits.

    (A) Marinoan (ca. 645 to 635 Ma) and (B) Sturtian (717 to 659 Ma) deposits (28, 33). Yellow dots indicate regional-scale deposits of glacial and/or periglacial origin. Red dots indicate glacial-periglacial deposits with associated sedimentary Fe oxide formation (160, 400). Black stars in yellow dots indicate occurrence of authigenic and/or seafloor barite (BaSO4) in the postglacial cap dolostone (131). Areas lacking glacial deposits, for example, northeastern North America and central Europe, simply lack Cryogenian sedimentary records.

  • Fig. 5 Cryogenian paleogeography and the breakup of Rodinia.

    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).

  • Fig. 6 Sediment accumulation rates during Cryogenian and younger glacial epochs (169).

    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).

  • Fig. 7 Images of the Cryogenian sedimentary record.

    (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 CO2 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) (486488). (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 (Fe2O3 + SiO2) 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 (BaSO4) 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 CO2, the size of the seawater sulfate reservoir, the elimination of the atmospheric Δ17O anomaly, and the time scale for cap dolostone sedimentation (91).

  • Fig. 8 Snowball Earth intermodel comparison (81).

    Colors assigned each GCM as indicated in (B). Solid lines, CO2 = 0.1 mbar; dashed lines, CO2 = 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 CO2 and H2O. 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 CO2 level (80, 82, 228).

  • Fig. 9 More Snowball Earth intermodel comparisons (81).

    See Fig. 8 caption for the prescribed conditions. January mean mass Eulerian stream function for models SP-CAM, LMDz, and ECHAM at (A) CO2 = 0.1 mbar and (B) CO2 = 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 × 109 kg s−1. Maximum mass stream functions in January increase by a factor of ~1.5 between 0.1 and 100 mbar of CO2. ECHAM is unstable at CO2 = 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 × 109 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) CO2 = 0.1 mbar and (F) CO2 = 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 CO2 but is muted in FOAM due to cold surface temperatures (Fig. 8, A and B). LMDz exhibits grid-scale noise in (F).

  • Fig. 10 Ice-sheet recession on Snowball Earth with rising CO2 (100).

    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 CO2 = 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 CO2.

  • Fig. 11 Sensitivity of ice sheets on Snowball Earth to precession-like forcing (100).

    Same coupled atmosphere–ice-sheet model as in Fig. 10, with CO2 = 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.

  • Fig. 12 Cryoconite distribution on glacial ice.

    (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).

  • Fig. 13 Modern Antarctic analogs for Cryogenian sublimating ice surfaces.

    (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).

  • Fig. 14 Meltwater cycles in polar and low-latitude cryoconite holes.

    (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 O2 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 CO2 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 (293296).

  • Fig. 15 Non–steady-state and steady-state hydrologic cycling on Snowball Earth, with and without continents.

    (A) Non–steady-state Snowball aquaplanet on which low-albedo marine ice (497) outcrops in the sublimation zone. Sublimation of marine ice (magenta arrow) is not balanced by a return flux to the seawater–marine ice subsystem (334). (B) Steady-state Snowball aquaplanet on which only meteoric ice is exposed. (C) Non–steady-state Snowball Earth with continents. Meteoric ice-sheet meltwater enters the ocean at ice grounding lines (magenta arrows) but is not balanced by a return flux to the atmosphere–meteoric ice subsystem, because only meteoric ice is exposed. (D) Steady-state Snowball Earth with continents. Ice-sheet meltwater injected into the ocean is balanced by sublimation of outcropping marine ice. Cryoconite meltwater flushing (Fig. 18B) functions essentially like a grounded ice sheet, fluxing meteoric water into the subglacial ocean.

  • Fig. 16 2D sea-glacier flow model in a Cryogenian paleogeography.

    Steady-state ice thickness (scale bar in meters) and velocity field (arrows, in meters per year, with every fourth velocity vector shown) with surface temperatures smoothly fitted to NCAR GCM results assuming a surface albedo of 0.6 and CO2 = 0.1 mbar (A) and 100 mbar (B) (291). Paleogeography from the study by Li et al. (179). Tropical ice thickness changes as CO2 rises are minimal, because surface warming softens the ice, reducing the meridional thickness gradient. (C and D) Log10 of the effective viscosity (291) with the same albedo and CO2 forcings as (A) and (B), respectively. The thinnest, softest, and fastest-flowing ice occurs in meridional sounds and gulfs.

  • 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.

  • 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 O2 is added to the atmosphere. SL, sea level.

  • Fig. 19 Cryogenian glacial deposits and cap carbonates.

    (A) Volcanic ash layer (arrow) in Marinoan marine periglacial Ghaub Formation on the foreslope of the Otavi Group carbonate platform, Fransfontein, Namibia. Detrital carbonate host includes suspension fallout (tan) and turbidites (gray). The pen is 15 cm long. The reddish color is Fe stain related to the ash layer. Such Marinoan ash layers are cited as evidence for open water (83), but the ash could have fallen anywhere on a sea glacier and could have been advected to the sublimation zone (Fig. 12C), where it would eventually be flushed through a moulin into the subglacial ocean (Fig. 18B). The flushing process would concentrate the ash spatially as many interconnected ponds may drain through a single moulin. The potential for dispersal by currents in the subglacial ocean is a fate that is shared by ash falling into an ice-free ocean. (B) Sturtian synglacial iron formation with ice-rafted dropstone (extrabasinal quartz monzonite) in the uppermost Sayunei Formation (Rapitan Group), near Hayhook Lake, Mackenzie Mountains, Northwest Territories, Canada. The pen is 15 cm long. The iron formation occurs directly beneath a kilometer-thick ice grounding-zone diamictite complex (Shezal Formation) and is inferred to have been deposited in front of an advancing tidewater ice margin. (C) Sturtian cap dolostone [basal Tapley Hill Formation (TH)] with low-angle cross-bedding sharply overlies syndeglacial siltstone [Lyndhurst/Wilyerpa Formation (LH)] bearing ice-rafted dropstones (arrow), Kingsmill Creek, near Tillite Gorge, Arkaroola Wilderness Sanctuary, Northern Flinders Ranges, South Australia. The 1- to 2-m-thick siltstone is underlain by ~1.5 km of synglacial boulder diamictite (E) of the Sturtian Merinjina/Bolla Bollana Formation (Umbaratana Group). The hammer handle is 33 cm long. The shallow-water cap dolostone is unusually well developed for a Sturtian cap-carbonate sequence (476). (D) A typical Sturtian cap carbonate—micritic organic-rich limestone with graded calcilutite turbidites, basal Twitya Formation (Windermere Supergroup), Gayna River, Mackenzie Mountains, Northwest Territories, Canada. The coin is 2 cm in diameter. (E) Massive Sturtian diamictite of the Merinjina/Bolla Bollana Formation, Tillite Gorge, Arkaroola Wilderness Sanctuary, Northern Flinders Ranges, South Australia. A ~1.5-km-thick ice grounding-zone deposit, composed of massive and stratified diamictites and conglomerate (497, 498), is draped by deglacial siltstone and a shallow-water Sturtian cap dolostone (C). (F) Sturtian glaciomarine sequence (Rapitan Group) and cap limestone [basal Twitya Formation (Fm)] near Stoneknife River (64°41.822′N, 129°53.629′W), Mackenzie Mountains, Northwest Territories, Canada. The 114-m-thick Rapitan Group comprises three massive (D1, D2, and D3) and two stratified (S) diamictite units. It disconformably overlies carbonates of the Little Dal and Coppercap formations. The sharp-based, 40-m-thick cap limestone features hummocky cross-bedding basally, indicating accumulation above storm wave base. The cap limestone is gradationally overlain by dark gray shale (maximum flooding) and a siltstone-dominated highstand system tract hundreds of meters thick. Paleomagnetic data (33, 127) indicate a subtropical paleolatitude for this location at the Sturtian glacial onset, consistent with the carbonate-rich pre- and post-Sturtian succession (148).

  • Fig. 20 Snowball Earth ocean dynamics.

    Results of a 2D (depth and meridian) ocean model (MITgcm) coupled to a 1D (meridian) ice-flow model (390, 391). Mid-ocean ridge (MOR) has an associated geothermal anomaly, and its elevation is corrected for glacial eustatic lowering. Depth is relative to glacial sea level. (A) Temperature, (B) salinity, (C) MOC stream function, and (D) zonal velocity, all with the MOR at 20°N. The total ranges of temperature and salinity are small. Strong low-latitude MOC (35 sverdrup) compares with the present high-latitude North Atlantic MOC (~20 sverdrup). (E) As in (D), but with the MOR removed and geothermal anomaly retained. (F) As in (D), but with the MOR located at the equator. (G) As in (F), but with the MOR removed and geothermal anomaly retained. At shallow depth beneath the sea glacier, zonal flow is directed eastward when the geothermal flux is symmetrical about the equator (F and G). When the geothermal flux is antisymmetrical (D and E), a westward-flowing jet occurs at shallow depth and low latitude in the warmer hemisphere, whereas an eastward-flowing jet occurs in the colder hemisphere. Zonal flow pattern is governed by the geothermal field, not by MOR topography. (H) Results of 3D high-resolution sector ocean model showing time-dependent turbulent eddy field. Snapshot of zonal velocity field at 125 m below the ice (1150 m below the surface). The white oval indicates an idealized continent. (I) Freezing rate (negative values imply melting) at the ice base (red) and prescribed sublimation and deposition rates at the ice surface (blue) as a function of latitude in the 2D model. The comparison shows that basal freezing/melting contributes as much as if not more than surface deposition/sublimation to sea-glacier thickness.

  • Fig. 21 Snowball Earth ocean circulation, 3D results with Marinoan paleocontinents.

    (A to F) Zonal (A and D), meridional (B and E), and vertical velocity (C and F) fields at depths of −1.1 km (A to C) and −2.9 km (D to F) relative to the ice surface (391). Ice thicknesses are as given in Fig. 20 (A to D). Scale bars are in centimeters per second (A, B, D, and E) and 10−6 m s−1 (C and F). Positive velocities are westward (A and D), northward (B and E), and upward (C and F). Zonal and meridional velocities change sign with depth, but vertical velocities do not. Strong vertical mixing (C and F) in the zone where cryoconite flushing occurs (Fig. 18B) expedites abyssal sedimentation of finely suspended material. Zonal shallow jets (A) direct flushed cryoconite toward western or eastern ocean margins, in response to excess geothermal input (as in Fig. 20C) from the northern hemisphere.

  • Fig. 22 Depositional model for synglacial iron formation.

    Fe isotope and redox proxy data suggest that Sturtian iron formations (Figs. 4B and 5B) were deposited at redoxclines where ferruginous waters upwelled into oxygenated meltwater (160, 412). Sedimentary facies associations (Figs. 7E and 19B) and stratigraphic relations imply that iron formation (IF) was deposited near ice grounding lines in semirestricted basins (for example, rift basins, glacial fjords, and overdeeps) (220, 400, 404, 409, 415). Ice-rafted debris (IRD) is concentrated proximal to the ice grounding line because shelf ice moves seaward more slowly than free-floating icebergs.

  • Fig. 23 Global distribution of sharp-based Marinoan postglacial cap dolostones.

    Distribution of Marinoan cap dolostones over 110° paleolatitude, from South China at 45°N (H) to West Africa at 65°S (E) (Fig. 5A), records an anomalous state of carbonate oversaturation in the surface ocean despite extreme CO2 acidification in the Snowball aftermath (73, 74, 8491). Other Marinoan cap dolostones are imaged in Figs. 7C and 25B. PG, preglacial carbonate strata; G, glacial-periglacial deposits; CD, cap dolostone. Arrows mark sharp basal contacts. (A) Ravensthroat Formation (CD) on Stelfox Member, Ice Brook Formation (G), and Keele Formation (PG), Arctic Red River, Mackenzie Mountains, Northwest Territories, Canada (409, 416, 499). (B) Basal Canyon Formation (CD) on Storeelv Formation (G), Tillitekløft, Kap Weber, Kejser Franz Joseph Fjord, East Greenland (139, 334). Resistant layers of quartz sandstone within glacial diamictite were eroded from unlithified preglacial marine sandstones. Photo courtesy of E. W. Domack. (C) Basal Zhamoketi Formation (CD) on Tereeken Formation (G), Yukkengol, Quruqtagh, Xinjiang, China (500). A lens cap is used for scale. Photo courtesy of S.-H. Xiao. (D) Basal Ol Formation (CD) on Khongor Formation (G) and Taishir Formation (PG), Khongoryn, Zavkhan, western Mongolia (143, 501). (E) Amogjar unit (CD) on Jbéliat Group (G), Atar, Adrar, Taoudeni Basin, Mauritania (161). (F) Nuccaleena Formation (CD) on Elatina Formation (G), Elatina Creek, Central Flinders Ranges, South Australia (427). (G) Sentinel Peak Member (CD) of the Noonday Dolomite on Wildrose Member (G) of the Kingston Peak Formation, Goller Wash, Panamint Range, Death Valley area, eastern California, USA (469, 470). (H) Basal Doushantuo Formation (CD) on Nantuo Formation (G), Huajipo section, Yangtze Gorges, northern Hunan, South China (417). Tianzhushania spinosa arrow indicates the first appearance of chert nodules preserving diapause cysts containing eggs and embryos of stem-group metazoans (502, 503). The yellow oval marks a volcanic ash layer yielding zircons dated at 635.2 ± 0.5 Ma (57). (I) Cumberland Creek Dolostone (CD) on Cottons Breccia (G), Grimes Creek, King Island, Tasmania (504). Igneous zircons from the topmost meter of the Cottons Breccia are dated at 636.41 ± 0.45 Ma (59).

  • Fig. 24 Lithostratigraphic comparison of idealized Sturtian and Marinoan cap-carbonate sequences.

    Closed arrows indicate the upward-deepening TST, MF indicates the maximum flooding stage, and open arrows indicate the upward-shallowing HST. TSTs are thin or absent in most Sturtian examples (but see Fig. 19C). Vertical bar indicates limestone (dark gray) or dolostone (light gray). “Cap dolostone” was deposited in surface water (above prevailing wave base) diachronously during shelf flooding, based on sedimentary structures indicated on the right. Numbers in blue are keyed to images in the accompanying figures. Tubestone stromatolite (441), giant wave ripples (443), and seafloor barite (91) are rare or absent outside of Marinoan cap dolostones, and benthic aragonite fans are developed to the same degree only in carbonates >1.5 Ga (423, 505). Sheet-crack cements (Fig. 25E) substitute for tubestone stromatolite in downslope areas (425). Relatively condensed Sturtian TSTs may reflect more deeply subsided continental margins because of the greater duration of the Sturtian cryochron (Fig. 2). Expanded Sturtian TSTs would be shifted landward, where net subsidence was less, but the record was more susceptible to later erosion.

  • Fig. 25 Distinctive features of Sturtian and Marinoan cap carbonates.

    (A) Sharp sedimentary contact between Sturtian glaciomarine Maikhan Ul Formation with ice-rafted dropstones (arrow) and overlying organic-rich and fossiliferous cap limestone (basal Taishir Formation), Zavkhan basin, western Mongolia (63, 143). Co-author F.A.M. points to the contact. (B) Sharp sedimentary contact between Marinoan glaciomarine Ghaub Formation (Bethanis Member) with ice-rafted dropstones and overlying cap dolostone (Keilberg Member of the Maieberg Formation) on the foreslope of the Otavi carbonate platform, northern Namibia (97). Co-author G.P.H. points to the contact. (C) Microbialaminite with roll-up structures in Sturtian cap carbonate (middle Rasthof Formation), northern Namibia. The coin is 2 cm in diameter. Roll-ups are associated with neptunian dikes and indicate that biomats were cohesive and pliable (475, 478), but the metabolic basis for their growth, inferentially below the photic zone, remains undetermined. (D) Macropeloids and low-angle cross-stratification in characteristically organic-poor Marinoan cap dolostone, Keilberg Member of the Maieberg Formation, northern Namibia (97). Sorted peloids and low-angle cross-stratification are characteristic of cap dolostones and indicate sedimentation at depths above fair-weather wave base. Coarse grain size suggests that macropeloids were less dense than pure carbonate when deposited. (E) Sheet-crack cements composed of fibrous isopachous dolomite (tinted orange by desert varnish where selectively silicified) occur regionally in the basal 2 m of the Marinoan cap dolostone in downslope settings in Namibia (425). The pen (circled) is 15 cm long. Sheet cracks and cements imply pore-fluid overpressure and an alkalinity pump, respectively, at shallow depth beneath the sediment-water interface at the onset of cap carbonate sedimentation. Overpressure could be related to rapid base-level fall due to the gravitational effect on the ocean of the loss of nearby ice sheets (419). (F) Upward-expanding mound of tubestone stromatolite (Figs. 24 and 26B) characterized internally by geoplumb (paleovertical) tubes (arrow) that were filled by carbonate mud as the mound grew. Mound margin (dotted line) is flanked by mechanically bedded dolopelarenite (lower left), Marinoan cap dolostone (Keilberg Member), northern Namibia (97). (G) Giant wave ripple (Fig. 24) in the Marinoan cap dolostone (Nuccaleena Formation) at Elatina Creek, central Flinders Ranges, South Australia. Note gradual amplification at the level of the hammer (circled, 35 cm long) and correlation of overlying layers a and b. Such unusually large and steep wave ripples occur in Marinoan cap dolostones on several paleocontinents, and the regional and global mean paleo-orientations of their crestlines are meridional (31, 443). Single wave trains aggrade as much as 1.4 m vertically, implying steady growth under the influence of long-period waves during Marinoan deglaciation (443). Intraclasts and composite grains are typically absent, suggesting that the characteristic steepness of the ripples is not related to early lithification [contra Lamb et al. (506)]. (H) Calcitized crystal fans (gray) formed as prismatic aragonite seafloor cement (Figs. 24 and 26C) in the Hayhook Formation stratigraphically above the Marinoan cap dolostone in the Mackenzie Mountains, Northwest Territories, Canada (409, 416). Growth of seafloor cement was coeval with sedimentation of lime mud (pink) from suspension. Seafloor aragonite precipitates up to 90 m thick occur in Marinoan cap limestones and mark the return of a depositional style last common before ~1.5 Ga (421, 423).

  • Fig. 26 Three-stage paleoceanographic model for Marinoan cap-carbonate sequences.

    (A) During the Snowball cryochron, net base-level fall (glacial eustasy > glacial isostasy + ice gravity) causes falling-stand and ice grounding-zone wedges to form on the continental slope, whereas lodgement till is deposited locally above sea level. Salinity of seawater may double, as a result of freshwater sequestration in ice sheets and sea glacier, and the freezing point is lowered accordingly. Glacial action supplies carbonate eroded from the platform to the ocean as suspended load in meltwater plume and grounding-zone wedge. (B) During deglaciation, meltwater production and surface warming create stable density stratification (387). The kilometer-thick meltwater lid (Glacial Lake Harland) includes an oxygenated mixed layer and intermediate water that is at least regionally anoxic, presumably due to respiration of surface organic export production. Marinoan cap dolostone was deposited diachronously from the meltwater mixed layer as base level rose and the glacierized landscape was flooded. The time scale for base-level rise may exceed the ice-sheet melting time scale because of whole-ocean warming and thermal expansion as destratification occurs (387) and over the peripheral bulges of former ice sheets, which collapse on the time scale of isostatic adjustment (IA) (419). (C) After ice sheets have melted, the meltwater lid mixes with the chemically evolved Snowball brine, which finally warms (387). The HST of cap-carbonate sequences fills in areas where net accommodation remains after GIA and hydro-isostatic adjustments. Landward progradation is inferred from inclined surfaces of constant carbonate δ13C (97).

  • Fig. 27 Relative sea-level (base-level) changes at Snowball terminations.

    Predicted base-level changes for a globally synchronous Marinoan deglaciation, during which ice sheets decrease in thickness linearly from a glacial maximum (A) to zero thickness in 2 ky while maintaining equilibrium profiles (419). Predictions incorporate the effects of glacial eustasy (global mean sea level), ocean–ice-sheet gravitational attraction, and GIA and hydro-isostatic adjustment on a rotating planet. However, they do not incorporate the effects of ocean warming and thermal expansion (387). (A) Cryogenian paleogeography (419) assumes maximum (165, 239) rather than fractional (100) ice-sheet volume at the termination. Scale bars are in meters relative to sea level. (B) Net base-level changes computed over the 2-ky deglaciation phase. Negative change is base-level fall. Scale bar is in meters. (C) Net base-level changes computed over the 10-ky interval following the end of deglaciation. The scale bar is in meters. Note the time dependency of base-level changes and their sensitivity to location across the ocean-continent interface. Preterminal ice-sheet recession (100) changes the spatial distribution of ice-gravity and isostatic effects and reduces their amplitudes as well as that of eustasy.

  • Fig. 28 Bar graph of peer-reviewed papers on Cryogenian glaciation by discipline and year, 1982 to 2016.

    Papers are assigned to one of four disciplinary categories. Note the relative growth of geophysical and geochemical papers after 1996 and geobiological papers after 2002. From the 1870s through the 1980s, research on “eo-Cambrian” glaciation was almost exclusively geological (55). The 70 chapters by various authors in the work of Arnaud et al. (175) are not tallied here. Forward modeling papers (in all categories) account for only ~25% of the current total.


  • Table 1 U-Pb and Re-Os geochronological constraints on Cryogenian glacial onsets and terminations.

    CA, chemical abrasion; ID-TIMS, isotope-dilution and thermal-ionization mass spectrometry; SIMS, secondary-ion mass spectrometry.

    PaleocontinentAge (Ma)Method*Reference
    Marinoan deglaciation/cap carbonate: 636.0 to 634.7 Ma
    Laurentia>632.3 ± 5.9Re-Os(63)
    South China635.2 ± 0.5U-Pb ID-TIMS(57)
    Southern Australia636.41 ± 0.45U-Pb CA-ID-TIMS(59)
    Swakop635.21 ± 0.59/0.61/0.92U-Pb CA-ID-TIMS(56, 83)
    Marinoan glacial onset: 649.9 to 639.0 Ma
    Congo>639.29 ± 0.26/0.31/0.75U-Pb CA-ID-TIMS(83)
    Southern Australia<645.1 ± 4.8Re-Os(137)
    South China<654.2 ± 2.7U-Pb SIMS(134)
    South China<654.5 ± 3.8U-Pb SIMS(58)
    Sturtian deglaciation/cap carbonate: 659.3 to 658.5 Ma
    Southern Australia>657.2 ± 2.4Re-Os(137)
    Tuva-Mongolia659.0 ± 4.5Re-Os(63)
    Southern Australia<659.7 ± 5.3U-Pb SIMS(366)
    Laurentia662.4 ± 3.9Re-Os(60)
    South China>662.7 ± 6.2U-Pb SIMS(65)
    Sturtian glacial onset: 717.5 to 716.3 Ma
    Oman>713.7 ± 0.5U-Pb ID-TIMS(365)
    South China<714.6 ± 5.2U-Pb SIMS(64)
    South China<715.9 ± 2.8U-Pb SIMS(62)
    South China<716.1 ± 3.4U-Pb SIMS(62)
    Laurentia>716.5 ± 0.2U-Pb CA-ID-TIMS(32)
    Laurentia<717.4 ± 0.1U-Pb CA-ID-TIMs(32)
    Laurentia<719.47 ± 0.29U-Pb CA-ID-TIMS(61)

    * Re-Os isochron ages from sedimentary organic matter. Errors are quoted at the 2σ level of uncertainty. Where multiple uncertainties are given, they represent analytical/analytical + tracer solution/analytical + tracer solution + decay-constant uncertainties.