Kirschvink (1992) highlighted three implications of his theory which "might lend themselves to geological tests". First, "glacial units should be more or less synchronous Harland, 1964)". Second, "strata from widely separate areas which preserve a record of these [rapid] climatic fluctuations [between 'ice house' and 'greenhouse' states] might bear an overall similarity in lithologic character, which would be a result of the global scale of the climatic fluctuations". Third, "the presence of floating pack ice should reduce evaporation, act to decouple oceanic currents from wind patterns and, by inhibiting oceanic to atmospheric exchange of O₂, would enable the oceanic bottom waters to stagnate and become anoxic. Over time, ferrous iron generated at the mid-oceanic ridges or leached from the bottom sediments would build up in solution and, when circulation became reestablished toward the end of the glacial period, the iron would oxidize to form a 'last-gasp' blanket of banded iron formation in upwelling areas." Here is how these tests turned out.

Test 1: Carbon isotopes
Direct radiometric testing of predicted glacial synchroneity has not proved possible (Table 1) because of an apparent dearth of suitable material for dating (Evans, 2000). Low diversity and turnover rates of microfloral and protistan remains pose severe limits on biostratigraphic resolution until after the major Neoproterozoic glacial events (Knoll, 2000). Regardless, a sea change in favor of synchronous glaciation occurred in the 1990s, and carbon isotope stratigraphy was principally responsible.

The history of carbon isotopic variation (expressed as δ¹³C, the difference in per mil of ¹³C/¹²C with respect to standard VPDB) in marine carbonates and covarying organic matter is a robust, but ambiguous, record of biogeochemical cycling (Summons and Hayes, 1992). Despite the fact that significant spatial heterogeneity in δ¹³C is dynamically maintained within the ocean (Kroopnick, 1985), large secular changes in δ¹³C at individual sites are globally correlated (Veizer et al., 1999; Saltzman et al., 2000), reflecting the residence time of carbon ~150 times the ocean mixing time (Kump and Arthur, 1999). The δ¹³C of carbonate rocks is relatively immune from large diagenetic changes because pore fluids are effectively buffered by the large rock carbon reservoir (Veizer et al., 1999).



The range of secular δ¹³C variability of >10‰ in the late Neoproterozoic (Jacobsen and Kaufman, 1999; Walter et al., 2000) is markedly greater than for the Phanerozoic (Veizer et al., 1999) or the 1.5 billion years before the late Neoproterozoic (Brasier and Lindsay, 1998; Kah et al., 2001). The largest Neoproterozoic δ¹³C anomalies consistently bracket LNGD (Fig. 8,9), with enriched values (δ¹³Ccarb>5‰) observed prior to glaciation and depleted values (δ¹³Ccarb<0‰) afterwards. A steep decline in δ¹³C values of 10-15‰ precedes any physical evidence of glaciation or sea-level fall in many areas (Narbonne and Aitken, 1995; Hoffman et al., 1998b; Brasier and Shields, 2000; McKirdy et al., 2001; Xiao et al., 2001; Halverson et al., 2002). Cap carbonates are universally depleted, in some cases approaching mantle values of –6±1‰ (Knoll and Walter, 1992; Kaufman and Knoll, 1995; Kennedy, 1996; Kaufman et al., 1997; Hoffman et al., 1998b; Prave, 1999; Brasier and Shields, 2000; Walter et al., 2000; James et al., 2001). Although interpretations of these remarkable trans-glacial isotopic shifts differ widely (see below), there is consensus that they are supra-regional in nature. Even those opposed to the snowball hypothesis (Kennedy et al., 1998, 2001a,b; Prave, 1999a,b; Condon and Prave, 2000) agree on the synchroneity of glacial events in widely separate areas (Africa, Australia, Canada, Europe) and base global correlation schemes upon them.



The conventional explanation for Neoproterozoic carbon isotopic oscillations and cap carbonates is oceanographic (Knoll et al., 1986; Kaufman et al., 1991; Derry et al., 1992; Grotzinger and Knoll, 1995; Knoll et al., 1996; Kaufman et al., 1997; Grotzinger and James, 2000). The idea is roughly the following. The ocean was physically stratified for long periods pending glaciation, but switched to a turnover mode during and particularly after glacial events. In stratified mode, the biological pump (descent of organic particles from the surface) drove the δ¹³C of dissolved inorganic carbon (DIC) in the surface waters to higher values. Partial remineralization of the organic "rain" in conjunction with bacterial sulfate reduction in the water column produced deepwater laden with 13C-depleted DIC. With prolonged stratification, the contrast in δ¹³C values between the surface and deep water DIC reservoirs (2-3 per mil in the present ocean, Kroopnick, 1985) increased to >10 per mil, consistent with the pre-glacial high δ¹³C values (Fig. 8). When turnover began, anoxic, alkalinity-laden, deep water upwelled to the surface, releasing CO₂ and precipitating cap carbonates with low δ¹³C values. Like the snowball hypothesis, the ocean "turnover" model (Grotzinger and Knoll, 1995; Knoll et al., 1996) seeks to explain many of the salient observations, but the basic premise is faulty. First, it is physically difficult if not impossible to shut down the overturning circulation for periods >1.5 kyrs (Zhang et al., 2001). Second, organic productivity would crash if the nutrient-rich upwelling flux were removed (Hotinski et al., 2001). Consequently, stratification would not raise the δ¹³C in surface waters to 5-10‰, as is observed. Third, cap carbonates formed during post-glacial sea-level rise (see below), when the injection of glacial meltwater would suppress, not enhance, ocean turnover. Kennedy (1996) attributes cap carbonates and their low δ¹³C values simply to post-glacial sea-level rise, but his model also depends on an 8-10‰ difference in δ¹³C across the thermocline, compared with ~2.5‰ for Mesozoic-Cenozoic oceans.

In 1998, we invoked the snowball hypothesis, then in eclipse, as a possible explanation for the δ¹³C profiles (Fig. 8,9) of carbonates bounding glacial horizons in northern Namibia (Table 1). Hoffman et al. (1998b) gave three, mutually compatible explanations for the ¹²C enrichment observed in cap carbonates. First, organic productivity on snowball Earth would be greatly reduced for 106-107 years, which should cause δ¹³C values to shift towards the hydrothermal CO₂ input (–6±1‰ VPDB, Des Marais and Moore, 1984) at mid-ocean ridges (Kump, 1991). Second, if carbonate sedimentation rates globally were very high, due to the alkalinity flux from weathering in the greenhouse transient, the organic fraction of the burial flux would be negatively effected. Consequently, carbonate would have low δ¹³C values even if organic productivity had fully recovered from the snowball event (Hoffman and Schrag, 1999). Third, the transfer of atmospheric CO₂ to bicarbonate (via the weathering cycle) involves an ~8‰ isotopic fractionation favoring ¹³C. In transferring >98% of the built-up atmospheric CO₂ to the ocean and ultimately the sedimentary reservoir, the kinetic isotope effect (Rayleigh distillation) will drive the atmospheric CO₂ reservoir to ever lower δ¹³C values (Hoffman et al., 1998b). To the degree that atmospheric CO₂ remains a major source of DIC, this would be reflected by cap-carbonate δ¹³C values that decline with time (Fig. 8). However, the nadir in δ¹³C (Fig. 9) does necessarily correspond to the completion of atmospheric CO₂ drawdown, as Kennedy et al. (2001b) argue, but reflects the net effect of many factors, including those already mentioned, as well as methane hydrate destabilization during deglaciation (Kaufman et al., 1997; Kennedy et al., 2001a; but see Hoffman et al., 2002a) and dolomite-calcite isotopic equilibrium.

The swing to negative δ¹³C values observed beneath the glacial deposits in many regions (Fig. 8,9) must originate differently (Kennedy et al., 2001b; Schrag et al., 2002; Halverson et al., 2002) as they are separated in time from the cap-carbonate anomalies by the glacial periods which were of long duration (Hoffman et al., 1998a; Sohl et al., 1999). We will discuss the pre-glacial anomalies, both positive and negative (Fig. 8), and the larger problem of how snowball events originate, in the penultimate section of the paper.

Test 2: Cap carbonates
Most LNGD are capped by continuous layers (Fig. 3-5) of pure dolostone (and locally limestone), meters to tens of meters thick (Spencer and Spencer, 1972; Williams, 1979; Deynoux, 1980; Fairchild and Hambrey, 1984; Miller, 1985; Preiss, 1987; Aitken, 1991; Fairchild, 1993; Brookfield, 1994; Grotzinger and Knoll, 1995; Kennedy, 1996; Hoffman et al., 1998a; James et al., 2001; Corkeron and George; 2001). Their basal contacts are characteristically abrupt (Fig. 2c, 2f, 4a and 5a), with little evidence of reworking or significant hiatus (Norin, 1937; Williams, 1979; Eisbacher, 1981; Deynoux, 1982; Preiss, 1985; Hoffman et al., 1998a; Brasier et al., 2000). Cap dolostones are transgressive and they typically extend far beyond the confines of their antecedent glacial deposits, blanketing pre-glacial rocks disconformably (Wright et al., 1978; Preiss, 1985; Aitken, 1991; Hegenberger, 1993; Hoffmann and Prave, 1996). They range from thin, allodapic, deepwater deposits (Kennedy, 1996), through agitated shelfal facies (James et al., 2001), near-shore grainstones and stromatolites (Spencer and Spencer, 1972; Corkeron and George, 2001), to supratidal tepee beccias (Fig. 5g; Deynoux and Trompette, 1976). "Cap-carbonate sequences" (see below and Fig. 3) share a distinct set of unusual sedimentary structures (Cloud et al., 1974; Peryt et al., 1990; Aitken, 1991; Fairchild, 1993; Grotzinger and Knoll, 1995; Kennedy, 1996; Hoffman et al., 1998a; James et al., 2001), and they are consistently depleted in ¹³C relative to other Neoproterozoic carbonates (Knoll and Walter, 1992; Kaufman and Knoll, 1995; Kaufman et al., 1997; Hoffman et al., 1998b; Kennedy et al., 1998; Myrow and Kaufman, 1999; Brasier and Shields, 2000; Walter et al., 2000). They share certain features with Archean and some pre-Neoproterozoic carbonates (Grotzinger and Knoll, 1995; Grotzinger and James, 2000; Sumner, in press), but they are without parallel in the Phanerozoic. They have long been considered paradoxical because they suggest that an abrupt transformation from glacial to tropical conditions took place virtually everywhere (Spencer, 1971; Schermerhorn, 1974; Williams, 1979; Fairchild, 1993). Kirschvink (1992) did not foresee cap carbonates as consequential to snowball events—he was unaware of their existence (personal communication, 1999)—but they are a striking affirmation of the second testable implication (see above) of his theory. In cap carbonates, we see the "smoke", if not the "gun", of the snowball Earth.

"Cap dolostones" (Fig. 3) are the transgressive systems tracts (van Wagoner et al., 1988) of depositional sequences ("cap-carbonate sequences") associated with post-glacial sea-level rise. They typically grade across a marine flooding surface (Fig. 3) into deeper water limestone or shale, which gradually shoals upward through carbonate or siliciclastic lutites and arenites to an exposure surface 100s of meters above the base of the sequence (Fig. 4a,9). The great thickness and deep-water aspect of many cap-carbonate sequences (von der Borch et al., 1988; Narbonne and Aitken, 1995; Dalrymple and Narbonne, 1996; Hoffman et al., 1998b) is anomalous within their overall stratigraphic context (Fig. 9). This is commonly ascribed to post-glacial sea-level rise, but longterm tectonic subsidence is required to preserve the sequence from isostatically-driven erosion, because the glacio-eustatic cycle alone creates no permanent accommodation space. On the Otavi carbonate platform in northern Namibia, an estimated background tectonic subsidence rate equivalent to ~0.065 mm yr-1 of isostatically adjusted carbonate (Halverson et al., 2002) implies a timescale on the order of 5 million years to create the accommodation space occupied by the younger cap carbonate sequence (Fig. 9), which is 300-400 m thick. Hoffman et al. (1998a) suggested that that the snowball event occupied the vast majority of this time, when the average sedimentation rate on the platform was very low. By the end of the event, substantial net subsidence had occurred, although masked at that time by relative sea-level fall. Upon deglaciation, sea level rose and chemical and/or detrital sediments rapidly filled in the accommodation space provided, making up for "lost time", so to speak. High sedimentation rates in cap dolostones and overlying limestones are consistent with their low δ¹³C values (see above) and with their petrological characteristics (see below).

Petrologists being fascinated by the unusual have had a field day with cap carbonates (Peryt et al., 1990; Aitken, 1991; Fairchild, 1993; Grotzinger and Knoll, 1995; Kennedy, 1996; Hoffman et al., 1998b; James et al., 2001; Kennedy et al., 2001a; Sumner, in press; Corsetti and Kaufman, submitted). The Elatina, Ghaub, Ice Brook, Landrigan and Numees glacial deposits (Table 1) are directly overlain by pale, flinty, very pure, cap dolostones (Fig. 3), generally 3-30 m thick (Aitken, 1991; Kennedy, 1996; James et al., 2001). They are invariably well laminated and, at proper paleodepths, hummocky cross-laminated with laminae defined by sets of reverse-graded peloids (Fig. 4b)(Aitken, 1991; Kennedy, 1996; James et al., 2001). Structures referred to as "tepees" (Fig. 4c) occur in all the above but they have neither the geometry nor associations of conventional supratidal tepees (Kendall and Warren, 1987). In plan, they are linear and parallel, oriented subparallel to the regional paleoslope (Williams, 1979; Eisbacher, 1981; Aitken, 1991; Kennedy, 1996; James et al., 2001), not polygonal like conventional tepees (Fig. 5g), which result from lateral expansion due to the force of crystallization (Kendall and Warren, 1987). Inspection of their cuspate axial zones (Fig.; 4d) reveals that they formed by continuous aggradation in a highly oscillatory flow regime. Microbial bioherms (Fig. 5a-c) and biostromes (stromatolites) occur within some cap dolostones (Spencer and Spencer, 1972; Cloud et al., 1974; Wright et al., 1978; Walter and Bauld, 1983; Bertrand-Sarfati et al., 1997; Hoffman et al., 1998a; James et al., 2001; Corkeron and George, 2001). In Namibia and the North American Cordillera, stromatolites in cap dolostones of the Ghaub, Numees (and correlative Bildah), Wildrose, and Ice Brook glaciations (Table 1) are distinguished by internal, vertically oriented, tube-like or sheet-like pockets of infilled micritic sediment and cement (Cloud et al., 1974; Hegenberger, 1987, 1993; Hoffman et al., 1998a, 2002a). The tube-like variety (Fig. 5b) has been attributed to gas escape (Cloud et al.,1974; Hegenberger, 1987; Kennedy et al., 2001a), but close inspection of well-preserved examples (Fig. 5c) reveals a stromatolitic (microbial) origin (Hoffman et al., 1998a, 2002a; James et al., 2001). The micritic sediment gathered in shallow pits or gutters on the stromatolitic surface, which aggraded with remarkably little change in morphology (Fig. 5b). As a result, the sediment-filled pits and gutters became vertical tubes and sheets, respectively. The pitted stromatolite surfaces must have resembled giant golf balls (so designed to reduce drag). Cap dolostones have much more to reveal (Hoffman et al., 2002b), but the whole ensemble of structures suggests rapid aggradation on storm-dominated ramps.

Locally, cap dolostones are overlain by limestone buildups composed largely of sea-floor cements (Peryt et al., 1990; Aitken, 1991; Grotzinger and Knoll, 1995; Kennedy, 1996; Soffer, 1998; James et al., 2001). The cements consist of upward-fanning, prismatic crystals pseudomorphosed by calcite (Fig. 4f,5e). The square-tipped, pseudo-hexagonal habit of the crystals indicates that the primary phase was aragonite. Although individual crystal fans are small (<1 cm), they aggregate into reef-like masses (pseudo-stromatolites) that range from meters to decameters in scale (Fig. 5a,d). The cements did not precipitate slowly in the absence of particulate sediment, but grew rapidly in a race against micritic burial (Fig. 4f,5d,e). The cementstone buildups are up to 50 m thick, but eventually they give way to flaggy, allodapic, pink limestone with large-scale hummocky cross-bedding (Maieberg Formation, Namibia), or to black, organic-rich shale (Sheepbed Formation, Canada). The cements imply a high degree of carbonate oversaturation (Grotzinger and Knoll, 1995; Sumner, in press) and the dendritic-prismatic growth habit suggests rapid, diffusion-limited, growth rates (Lasaga, 1998). The presence of dissolved Fe(II) may have inhibited carbonate nucleation in anoxic waters below the mixed layer, thereby promoting the growth of crystal fans on the sea floor (Sumner and Grotzinger, 1996) while the micritic component nucleated in oxic surface waters. Deepwater anoxia is inevitable if the ocean was ice-covered for long periods (see below) and is consistent with dolomitization of cap dolostones on the sea floor (Baker and Kastner, 1981) and with the co-occurrence of sea-floor barite precipitates (Fig. 4e). The localization of these cementstone buildups at ancient shelf breaks (Soffer, 1998; James et al., 2001) suggests that their occurrence may be dictated by upwellings.

Differences among cap-carbonate sequences have elicited debate as to temporal versus spatial control (Kennedy et al., 1988; Brasier and Shields, 2000). Some examples (Fig. 3d,e) lack transgressive systems tracts and their isotopic profiles also appear base truncated—the descending phase of the δ¹³C anomaly is missing (Hoffman et al., 2001). This could result from undersaturation of the local water mass until after the post-glacial sea-level rise was complete, bearing in mind the large alkalinity input required to maintain saturation if atmospheric pCO₂ rose to high levels (Fig. 10) during the hypothesized snowball event. Base truncation appears more typical of "Sturtian" (Table 1) cap-carbonate sequences (Kennedy et al., 1998), and could mean that the glacial ocean was less well buffered by carbonate dissolution and/or that less carbonate was available for weathering during deglaciation (see below). While actively subsiding areas (Kimberley, Adelaide, Cordillera, Otavi) have thick (200-1200 m) post-glacial sequences (Plumb, 1981; von der Borch et al., 1988; Christie-Blick et al., 1995; Narbonne and Aitken, 1995; Hoffman et al., 1998a), the Jbéliat cap-carbonate sequence in the cratonic Taoudeni basin (Deynoux and Trompette, 1976) is highly condensed (Fig. 3,11). After isostatic adjustments, minimal accommodation space remained after post-glacial sea-level rise because the background tectonic subsidence rate was very low. The cap carbonate built rapidly to sea level, resulting in the formation of spectacular supratidal tepee structures (Fig. 5g) and associated, early diagenetic (phreatic), barite crusts (Fig. 5f). We speculate the latter may have formed in a meteoric-marine groundwater mixing zone beneath a surface exposed to the hypothesized transient greenhouse atmosphere. Where the sequence expands to fill paleovalleys (Fig. 11), the transgressive nature of the basal cap dolostone which drapes the paleotopography is apparent. The Jbéliat example is also interesting inasmuch as the glacial deposits (continental tillites and outwash) preceded the hypothesized snowball event, which is apparently represented by a regolith horizon (Fig. 11) with well-developed polygonal sand wedges (Deynoux, 1982, 1985). This horizon separates the glacial deposits from the cap dolostone, and marks a period when ablation exceeded precipitation in the region. Cap carbonates also thin basinward: the Maieberg Formation is 300-400 m thick on the Otavi platform (Fig. 9), but its equivalent (correlated isotopically) thins to <25 m within 30 km downslope from the shelf break (Hoffman et al., 1998b; Kennedy et al., 1998).

We are not burdened by an overabundance of explanations for cap carbonates. Fairchild (1993) suggested that they were precipitated from glacial meltwaters saturated with alkalinity due to interaction with carbonate-rich glacial debris and rock powder. This does not adequately explain the presence of cap carbonates in areas where the glacial deposits have little carbonate, for example the Kimberleys in Australia (Plumb, 1981). Grotzinger and Knoll (1995) proposed that cap carbonates record the turnover of a previously stratified ocean (see above), with upwelling of alkalinity-laden deepwater driving precipitation of 13C-depleted carbonates. This model dovetails with Kirschvink's (1992) post-glacial interpretation of Neoproterozoic BIF (Klein and Beukes, 1993), but the premise is questionable for reasons already given. Kennedy (1996) related cap carbonates to post-glacial flooding of continental shelves and seas, analogous to the "coral reef" hypothesis for the Quaternary (Berger, 1982), and Kennedy et al. (2001a) hypothesized that cap carbonates are byproducts of methane released from permafrost on continental shelves globally during post-glacial flooding. But methane oxidized aerobically in the water column produces CO₂, causing carbonate dissolution; only methane oxidized anaerobically in conjunction with bacterial sulfate reduction in pore waters yields bicarbonate (Boetius et al., 2000). If very large amounts of methane were released rapidly (Kennedy et al., 2001a), sulfate limitaton would likely prevent substantially increased rates of anaerobic methane oxidation. Moreover, the characteristic spatial variability in δ¹³Ccarb (-40-0‰) observed in Phanerozoic methane-seep carbonates (Kauffmann et al., 1996) is not found in cap carbonates. Tube-like structures purportedly conduits for methane release (Kennedy et al., 2001a) are strictly internal to cap carbonates (Fig. 5b): they do not tap the conjectured methane source rocks (Hoffman et al., 2002a). They occur only within microbial boundstones (Fig. 5c), even in cap dolostones deposited directly on metamorphic or granitic rocks (Cloud et al., 1974; Wright et al., 1978), unlikely to be prolific sources of methane.

Once we realized that the snowball hypothesis could explain the δ¹³C anomalies associated with LNGD, we took as a test that it should also account for cap carbonates. Recall the deglaciation of a snowball earth, sea ice and glaciers disappearing rapidly due to reverse ice-albedo feedback in an intense greenhouse with >10% CO₂ (Caldeira and Kasting, 1992). As surface temperatures soar, carbonic (and initially sulfuric) acid rain beats down on a landscape dominated by frost-shattered rock and rock powder from glacial action, not to mention millions of years' worth of unaltered volcanic material (Hoffman et al., 1998b; Hoffman and Schrag, 2000). Intense chemical (and physical) weathering ensues. Initially, where former carbonate platforms are exposed, carbonate weathering dominates. Later, after sea level has risen, silicate weathering becomes proportionally more important, and over time, draws down the atmospheric CO₂ until a new steady state is reached (Hoffman and Schrag, 2000). The time scale of the greenhyouse transient is unknown, but is conservatively estimated to be in the range 104-106 years [not <104 years as suggested by Kennedy et al. (2001b). The alkalinity flux driving cap carbonate deposition during post-glacial sea-level rise (i.e., below "maximum flooding" in Fig. 3, 9) came from carbonate weathering; silicate weathering is too slow, given a deglaciation time of <104 years (Kennedy et al., 2001b).

Large inputs of alkalinity are required to maintain seawater saturation if atmospheric pCO₂ greatly increased in a snowball event (Fig. 10). The hydrothermal dominance of the snowball ocean will drive down seawater pH unless it is buffered by dissolution of carbonate, supplied for example by glacial action. Buffering would enable cap carbonates to form immediately upon deglaciation. This is consistent with the inference (see above) that many cap carbonates formed during post-glacial transgression (Fig. 3) and with the absence of post-glacial karst beneath cap carbonates. Rapid deglaciation of a snowball Earth would emplace a growing pool of warm, oxic, low-density, weathering-dominated water above a cold, saline, anoxic, hydrothermally-dominated deepwater. The internal stratigraphy of cap carbonates (Fig. 3) may reflect the mixing of these contrasting water masses (Hoffman et al., 2002), as well as the shift from carbonate to silicate-dominated weathering.

Test 3: Iron formations
For Joe Kirschvink, iron is king of the elements. Iron formations are extensive deposits of sedimentary Fe(III) oxide and chert, not directly associated with volcanoes or hydrothermal vents. Common in the early stratigraphic record (Fig. 12), they virtually disappear after 1.85 Ga (Isley and Abbott, 1999). The only major reappearances of iron formation since that time (barring Clinton-type oolitic ironstones) occur strictly within LNGD (Fig. 2d), notably in Canada (Young, 1976; Yeo, 1981, 1986; Klein and Beukes, 1993), Brazil (Urban et al., 1992; Graf et al., 1994; Trompette et al., 1998), Namibia-South Africa (Martin, 1965), and Australia (Preiss, 1985; Lottermoser and Ashley, 2000).

The prerequisites for iron formation are three (Drever, 1974; Holland, 1984):

1. ocean anoxia to allow extensive transport of dissolved Fe(II);
2. an H₂S:Fe(II) flux ratio <2, so that not all Fe(II) is scavenged by sulfide (FeS2) burial, and
3. a local oxidant to drive the precipitation of an iron-formation precursor.

Before the first atmospheric oxidation event around 2.4 Ga (Karhu and Holland, 1996), deepwaters were anoxic and riverine sulfate input was low (Walker and Brimblecombe, 1985; Canfield et al., 2000; Farquhar et al., 2000). After 2.4 Ga, deepwater anoxia arguably persisted, but not until 1.8 Ga did oxidative weathering drive sulfate reduction in the ocean to levels sufficient to bury all Fe(II) as FeS₂ (Canfield, 1998). The reappearance of iron formation in LNGD, after a 1.5-billion-year absence, is consistent with ocean anoxia due to stagnation beneath near-global sea-ice cover (Kirschvink, 1992), and also with an attenuated flux of reduced sulfur due to the frozen surface (Canfield and Raiswell, 1999).

Kirschvink (1992) and his early supporters (Klein and Beukes, 1993) suggested that Neoproterozoic iron formations were deposited at the end of snowball events, when renewed thermohaline circulation would have driven Fe(II)-rich deepwater into the oxic zone in upwelling areas. This model has been criticized on the grounds that iron formation occurs stratigraphically within or below, rather than above, certain LNGD (Williams and Schmidt, 2000). However, glacial erosion at low latitudes will itself be strongly biased toward the end of a snowball event, when the hydrologic cycle has strengthened (Fig. 7). Alternatively, iron formation may have been deposited beneath the tropical ice pack if sea ice was thin (<20 m) or discontinuous (McKay, 2000; but see Warren et al., 2002). Oxygenic photosynthesis would drive precipitation, given an anoxic ocean rich in dissolved iron. The giant Rapitan iron formations in the northern Canadian Cordillera, for example, were deposited close to the equator (Park, 1997) and are overlain by <600 m of glacial diamictite and outwash (Yeo, 1981). [Young (1988) attributes the Rapitan iron formation to regional (but cryptic), rift-related, hydrothermal activity, but this is not supported by trace-element geochemistry (Klein and Beukes, 1993; Graf et al., 1994) and ignores the complete absence of volcanics in the Rapitan belt and the age distribution of iron formations globally (Fig. 12).]

Kennedy et al. (1998) pointed out that iron formation is more common in (but not limited to) "Sturtian" glacial deposits (Table 1). A speculative explanation for this observation is the hydrostatic pressure dependence of the H₂S:Fe(II) ratio in hydrothermal vent fluids at mid-ocean ridges (Kump and Seyfried, 2001). A large fall of sea level would result in a lowering of this ratio, favoring BIF deposition (Kump and Seyfried, 2001). Many of the "Sturtian" deposits occur in rift zones and they tend to be thicker, if more localized, than the "Varanger/Marinoan" deposits (Eisbacher, 1981; Miller, 1985; Preiss, 1985; Young and Gostin, 1991; Brookfield, 1994; Narbonne and Aitken, 1995; Aleinikoff et al., 1995; Hoffman and Schrag, 1999). Arguably, there was more topography at low latitudes globally, giving rise to more massive ice sheets. This would have caused a larger sea-level fall and, according to Kump and Seyfried (2001), a greater propensity for iron formation.