Geologic evidence for a mantle superplume event at 1.9 Ga



[1] Both preserved and restored areal distributions of Proterozoic marine intracratonic, passive margin, and platform sediments show a prominent peak at ∼1.9 Ga, indicating that shallow marine sediments were widespread on the continents and that sea level was high at this time. The chemical index of alteration in shales deposited at this time was high, suggesting warm climates, possibly due to enhanced CO2 levels in the atmosphere. High sea level and warms climate may also explain an abundance of black shale, banded iron formations, and shallow marine phosphate deposits and an increase in the number of occurrences and diversity of stromatolites in general and microdigitate stromatolites at 1.9 Ga. All of these observations are consistent with a 1.9-Ga superplume event. The occurrence of only a minor positive carbon isotope shift in marine carbonates at 1.9 Ga indicates that the relative rates of burial of organic and oxidized carbon remained about the same as at present. Slightly low 87Sr/86Sr isotopic ratios in seawater at 1.9 Ga reflect increased mantle input of Sr from the proposed superplume event, whereas higher ratios at 1.85–1.75 Ga may reflect increased input of continental Sr from a growing supercontinent. The first massive sulfate evaporites in the geologic record at 1.8–1.6 Ga follow the possible 1.9-Ga superplume event. This may reflect an increase in both oxidation state and carbonate deposition in the oceans as plume-related volcanism wanes.

1. Introduction

[2] A mantle superplume event occurs when many large plumes generated in the lower mantle rise to the base of the lithosphere in a short period of time (<100 Myr) [Condie, 1998]. This is similar to the usage in the work of Larson [1991a, 1991b], where he applied the term superplume to one or more large mantle plumes beneath the South Pacific during the mid-Cretaceous. On the basis of the distribution of ages of juvenile crust or plume-generated igneous rocks, Condie [1998] and Isley and Abbott [1999] have proposed several superplume events in the mantle during the last 3 Gyr. Condie [1998, 2000a] suggests that a superplume event is part of a superevent cycle, involving supercontinent formation and breakup. The formation of a supercontinent is associated with proposed superplume events at 1.9 and 2.7 Ga. In the model, superplume events are triggered by catastrophic collapse of descending slabs through the 660-km seismic discontinuity. Greff-Lefftz and Legros [1999] suggest that resonance between the outer core and solar tidal waves destabilizes the D″ layer above the core, leading to the episodic generation of numerous large superplumes. Whatever the triggering mechanism, superplume events occur over short periods of time and have major effects on other Earth systems.

[3] Larson [1991b] and Kerr [1998] showed that a mid-Cretaceous superplume event coincided with increases in surface temperature, deposition of black shales, a rise in sea level, elevated δ13C in seawater, and a decrease in the rate of magnetic reversals. The rise in sea level is inferred in two ways: through the sedimentary record of deposition on passive margins and in craton interiors and through the record of increased production of oceanic crust. Although the increased production rate of oceanic crust has been questioned by Hardebeck and Anderson [1996], the effects of a Cretaceous superplume event on platform and passive margin sedimentation are undisputed. Continental flooding and passive margin sedimentation increased greatly during the Cretaceous superplume event. Because episodes of increased continental flooding are relatively well preserved in the Precambrian stratigraphic record, we suggest that they can be used to test for the occurrence of Precambrian superplume events.

[4] The increase in global temperatures during the mid-Cretaceous superplume event is largely due to an increase in atmospheric CO2 rather than climate moderation by increased continental flooding [Barron et al., 1995]. The indirect result of this increase in global temperatures is an increase in continental weathering rates, which can be assessed by looking at the chemical index of alteration (CIA) of shales. If Precambrian superplume events were accompanied by increased atmospheric CO2, they should manifest themselves by increased CIA in shales. Because shales are relatively easy to preserve in cratonic sequences, they are likely to preserve a record of Precambrian global warming and cooling events.

[5] Because the two major superplume events proposed at 2.7 and 1.9 Ga were probably more intense than the mid-Cretaceous event [Condie, 1998], they also should have manifestations in the geologic record. In particular, large injections of CO2 into the atmosphere-ocean system should be reflected in sedimentary rocks. In this paper, we evaluate and discuss geologic features in both detrital and chemical sedimentary rocks that are consistent with, or support, a superplume event at 1.9 Ga.

2. Sea Level at 1.9 Ga

[6] Models of sea level through time are diverse and depend upon assumptions, such as crustal thickness and growth rate, freeboard (elevation of continents above mean sea level), and mantle cooling rate, none of which are well known [Taylor and McLennan, 1985; Galer, 1991; Gurnis, 1993; Eriksson, 1999; Eriksson et al., 1999]. During a superplume event, sea level should rise because of isostatic uplift and thermal erosion of oceanic lithosphere above mantle plume heads [Kerr, 1994; Lithgow-Bertelloni and Silver, 1998]. Formation of oceanic plateaus and enhanced activity of ocean ridges during a superplume event also may raise sea level by displacing seawater onto continental shelves [Larson, 1991b]. Supercontinent formation, on the other hand, should lower sea level as mantle upwellings develop beneath supercontinents [Gurnis, 1993; Condie et al., 2000].

[7] Although shallow marine sedimentary successions are widespread and well preserved in the Phanerozoic and Neoproterozoic, only small remnants of older successions remain in the geologic record. For this reason, it is not possible to use sequence stratigraphy to estimate sea level in rocks older than ∼800 Ma [Eriksson, 1999].

[8] However, remnants of Mesoproterozoic and Paleoproterozoic marine platform sediments are widespread on the continents, and hence their areal distribution as a function of time yields important insight into the relative elevation of sea level with time. In this study, areas of Proterozoic marine intracratonic, passive margin, and platform sediments (hereafter referred to as intracratonic sediments) are estimated from geologic maps ranging in scale from 1:106 to 1:12,000. In regions where numerous small outcrops of the same succession occur, a line is drawn around all of the outcrops, and they are scaled as one unit. One of the major problems encountered in estimating preserved areas of sediments is that of complex structure. Where sediments are strongly folded, area estimates are minimal. If structures are relatively simple (anticlines, synclines, etc.), attempts were made to estimate predeformational areal distributions by structural reconstructions. Our results, which we consider to represent minimal areas of Proterozoic and Archean intracratonic sediments, are summarized in Table 1. Also included in Table 1 are depositional ages and estimates of the error of these ages.

Table 1. Areas and Ages of Late Archean and Proterozoic Intracratonic Sedimentary Rocks
LocationPreserved Area, km2Restored Area, km2t, MaAge Error, MaReferences
Adelaide Basin, Australia840,0001,854,55880050Ambrose et al. [1981]; Myers et al. [1997]
Amadeus, Australia80,000152,25065026Powell et al. [1994]; Myers et al. [1997]
Officer Basin, Australia60,000114,18865050Zang [1995]
Brigham Group, Idaho13,90025,17660050Smith et al. [1994]
Pocatello, Idaho14,00025,35760050Link et al. [1993]; Smith et al. [1994]
East Greenland10,50019,983650100Fairchild and Hambrey [1995]
Nevada–California border area187,600333,12358020Kaufman and Knoll [1995]
Windemere Group, western Canada36,70077,11375040Jefferson and Young [1989]; Narbonne and Aitken [1995]
McKenzie Mountains, Northwest Territories, Canada29,20064,468800100Brookfield [1994]; Narbonne and Aitken [1995]
Grand Canyon Supergroup, Arizona26,40058,28680020Link et al. [1993]
Amundsen, Northwest Territories, Canada23,40051,663800100Kaufman and Knoll [1995]
Polarisbreen Group, Svalbard35,00063,39360020Fairchild and Hambrey [1984]
Varanger Peninsula, Norway41,70075,52860040Vidal and Moczydlowska [1995]
Russian Platform80,000152,25065050Vidal and Moczydlowska [1995]
Damara, Namibia1,125,0002,037,62160050Kukla and Stanistreet [1991]
Gariep Group, South Africa30,00062,72474515Tankard et al. [1982]
Nama Group, South Africa65,000114,28457025Gresse and Germs [1993]
Adrar, Mauritania112,000259,823850150Clauer et al. [1982]
Madina–Kouta, NW Africa300,000662,342800100Villeneuve [1989]
West Congo200,000380,62665045Kaufman and Knoll [1995]
Bambui Group, Brazil50,000105,05975025Marshak and Alkmim [1989]; Misi and Veizer [1998]
La Tinta Group, southern Brazil45,00089,98770050Cingolani and Bonhomme [1982]; Brito Neves and Cordani [1991]
Boqui–Jacadigo Groups, Bolivia20,00036,22460024Litherland et al. [1986]
Espinhaco Supergroup, eastern Brazil25,00049,99370020Marshak and Alkmim [1989]
Penganga Group, India40,00086,15377540Chaudhuri et al. [1989]
Raipur Group, India35,00094,193100075Kale and Phansalkar [1991]
Banzhou, southern China25,00067,281100080Li and McCulloch [1996]
Sinian, southern China100,000190,31365025Li and McCulloch [1996]
Bhander Group, India58,000156,0921000150Chakrabarti [1990]
Dengying Formation, southern China125,000249,96370035Zunyi et al. [1986]
Jixian Group, northern China20,00038,06365020Yun [1985]
Huainan Group, eastern China56,000106,57565025Zang and Walter [1992]
Kuruktag Suite, eastern Tarim25,00045,28060020Brookfield [1994]
Turukhansk, Siberia40,00097,503900100Sergeev et al. [1997]
Earaheedy Group, Western Australia100,000633,665186562Hynes and Gee [1986]; Myers et al. [1997]
Hatches Creek Group, central Australia50,000277,195173022Blake et al. [1987]; Myers et al. [1997]
Glengarry Group, Western Australia28,000183,682190085Myers et al. [1997]
Pine Creek basin, northern Australia80,000524,805190065Pagel et al. [1984]; Needham et al. [1988]
Reynolds Range, central Australia50,000297,086180065Dirks and Norman [1992]
Flinton Group, eastern Canada16,70049,621110030Moore and Thompson [1980]
Sibley Group, Canada34,750113,9981200100Ojakangas and Morey [1982]
Athabasca Group, northern Canada43,000231,4121700100Ramaekers [1981]; Kotzer and Kyser [1995]
Belt Supergroup, Montana82,600330,303140075Link et al. [1993]
Apache Group, Arizona42,000124,7951100100Link et al. [1993]
Sioux–Barboo Qtzt, north central United States144,500777,6521700125Greenberg and Brown [1984b]; Holm [1997]
Ortega Qtzt, New Mexico90,350486,234170080Soegaard and Eriksson [1989]
Mazatzal Group, Arizona16,70089,874170010Doe and Karlstrom [1991]
Hembrillo succession, New Mexico18,35098,754170030Alford [1987]
Ramah Group, Labrador17,800116,769190025Knight and Morgan [1981]
Marquette Supergroup, Minnesota34,200247,702200075Greenberg and Brown [1984a]
Animikie Group, Minnesota26,100189,0362000100Van Schmus [1976]; Sims et al. [1993]
Mistassini Group, eastern Canada13,30091,676195035Rivers [1997]
Kaniapiskau Group, eastern Canada25,000164,0021900150Rivers [1997]
Belcher Group, Canada123,800812,1361900150Donaldson and Ricketts [1979]; Hoffman [1988]; Chandler and Parrish [1989]
Goulburn–Bear Creek, Canada39,800277,069196050Campbell [1979]; Hoffman [1988]
Great Slave Supergroup, Canada58,380382,976190050Hoffman [1988]
Coronation Supergroup, Canada44,480291,792190040Hoffman [1988]
Baker Lake, Canada13,90086,781185050Aspler and Chiarenzelli [1997]
Amer Group, Canada610038,084185020Hoffman [1988]
Fox River, Canada26,270168,954188010Hoffman [1988]
Wollaston, Canada44,480277,700185025Hoffman [1988]
Yurmata Group, Russia60,000217,314130040Sergeev [1994]
Burzyan Group, Russia40,000194,977160050Sergeev [1994]
Ruzizian, Congo75,000543,2062000100Mendelsohn [1981]
Toro Group, Uganda25,000164,0021900100Goodwin [1991]
Mporokoso Group, Zambia50,000269,084170050Unrug [1984]
Jacobina Group, eastern Brazil10,50076,0492000100Ledru et al. [1997]
Roraima Supergroup, Guiana132,000710,3811700100Gibbs and Barron [1993]
Vaupes Supergroup, Guiana67,000295,803150060Gibbs and Barron [1993]
Rio Fresco, Brazil110,000666,6601820100Bonhomme et al. [1982]; Teixeira et al. [1989]
Beneficiente Group, northern Brazil113,750527,687155075Bonhomme et al. [1982]; Teixeira et al. [1989]
San Ignacio Group, Bolivia200,000848,7101460100Litherland et al. [1986]
Paranoa Group, India30,00089,139110050Fairchild et al. [1996]
Cuddapah Supergroup, southern India65,000349,809170050Meijerink et al. [1984]; Kale and Phansalkar [1991]
Kaladgi Group, India10,00053,8171700150Kale and Phansalkar [1991]
Purana basins, India90,000359,8941400100Kumar [1995]; Kale and Phansalkar [1991]
Vindhyan Supergroup, India90,000295,246120050Kumar [1995]; Kale and Phansalkar [1991]
Gangpur Group, India7,50038,988166525Kumar [1995]; Kale and Phansalkar [1991]
Sibao Group, southern China25,00095,143135040Li and McCulloch [1996]
Liaohe Group, northern China42,000289,504195040Zunyi et al. [1986]
Changchengian Formation, northern China25,550159,515185075Zunyi et al. [1986]
Derbina sequence, Siberia25,000115,9751550150Rosen et al. [1994]
Pahrump Group, California15,30044,570108020Wright et al. [1974]; Heaman and Grotzinger [1992]
St. Boniface Group, eastern Canada400012,738117030Rivers [1997]
Osler Group, eastern Canada23,00071,808115030Ojakangas and Morey [1982]
Mount Isa Supergroup, NE Australia50,000284,142175535Blake and Stewart [1992]
McArthur Early, northern Australia180,0001,017,859175010Rawlings [1999]
McArthur Intermediate, northern Australia45,000230,480165015Rawlings [1999]
McArthur Late, northern Australia80,000336,140145010Rawlings [1999]
Bangemall Early, Western Australia30,000148,420161515Collins and McDonald [1994]; Myers et al. [1997]
Bangemall Late, Western Australia30,000117,0321375100Collins and McDonald [1994]; Myers et al. [1997]
Birrindudu basin, northern Australia25,000118,294157020Sweet 1977]; Myers et al. [1997]
Yeneena Supergroup, NW Australia66,000216,5141200120Williams [1990]; Myers et al. [1997]
Victoria River basin, northern Australia35,000114,818120050Tyler and Griffin [1990]; Myers et al. [1997]
Hurwitz Group, Canada112,600994,1042200100Hoffman [1988]
Black Hills, South Dakota36,700272,470202520Hoffman [1988]
Snowy Pass Supergroup, Wyoming68,700669,646230075Hills and Houston [1979]; Karlstrom et al. [1983]
Huronian Supergroup, eastern Canada110,0001,072,213230075Fedo et al. [1997]; Rivers [1997]
Otish Group, eastern Canada16,100164,897235050Rivers [1997]
Lomagundi Group, Zimbabwe15,000119,9472100100Cahen and Snelling [1984]; Master [1991]
Piriwiri Group, Zimbabwe35,000253,496200075Cahen and Snelling [1984]; Master [1991]
Central Finland54,000526,3592300100Melezhik et al. [1997]
Minas Supergroup, Brazil10,800118,551242025Babinski et al. [1995]; Machado et al. [1996]
Prince Charles Land, Antarctica100,0001,311,819260050Sheraton et al. [1987]
Aravalli Group, India30,000374,540255050Verma and Greiling [1995]; Wiedenbeck et al. [1996]
Sandur Schist, India35,000415,860250050Manikyamba and Naqvi [1995]
Chitradurga Group, southern India25,000327,9552600100Argast and Donnelly [1986]
Hutuo Group, northern China30,000378,266256050Wang and Qiao [1984]
Shuanshanzi Group, northern China30,000239,8942100100Zunyi et al. [1986]
Wutai Group, northern China25,000297,0432500100Zunyi et al. [1986]
Hamersley Supergroup, Western Australia50,000594,085250080Trendall [1983]; Myers et al. [1997]
Turee Creek–Wyloo Groups, Western Australia35,000309,002220095Thorne and Seymour [1991]; Myers et al. [1997]
Yerrida Group, Western Australia75,000662,1472200110Pirajno et al. [1996]; Myers et al. [1997]
Hapschan Series, Siberia50,000538,0882400150Condie et al. [1991]
Udokan Series, Siberia45,000397,2882200100Rosen et al. [1994]
Dzheltulak Group, Siberia30,000252,068215050Rosen et al. [1994]
Zhongtiao Group, northern China15,000119,947210050Sun et al. [1990]
Transvaal Supergroup, South Africa105,0001,223,120248050Eriksson and Clendenin [1990]; Cheney and Twist [1986]
Witwatersrand Supergroup, South Africa37,500662,050290040Robb et al. [1990]; Winter [1994]
Pongola Supergroup, South Africa28,800528,993294060Cheney and Winter [1995]
Francevillian Supergroup, Gabon37,500299,868210065Ledru et al. [1989]
Jatulian successions, Finland25,000199,912210055Heiskanen [1991]

[9] The areal distribution of sediments from any particular tectonic setting varies with age because of the effects of erosion. Veizer and Jansen [1985] and Veizer [1988] have shown that the preservation of sediments decays exponentially with time and varies between tectonic settings. Phanerozoic intracratonic and platform sediments have a half-life of ∼350 Myr and an oblivion age (life expectancy) of ∼1600 Myr [Veizer, 1988]. Because there are numerous examples of intracratonic sediments preserved in the geologic record that are older than 1600 Ma, a longer half-life would seem appropriate for Proterozoic sediments. A more realistic half-life for Proterozoic intracratonic sediments is ∼700 Myr, with a corresponding oblivion age of ∼2500 Myr. In Table 1 (column 3) we have calculated the “restored” areas of the Precambrian sediments using the method of Veizer [1988] with a half-life of 700 Myr.

[10] The cumulative preserved area of Proterozoic marine intracratonic sediments is shown in histogram form as a function of age in Figure 1, and the restored sediment areas are shown in similar fashion in Figure 2. Both graphs show prominent peaks at 1.9–1.8 Ga and 800–600 Ma and a smaller peak at ∼2.7 Ga. This strongly suggests that shallow marine sediments were more widespread on the continents at these times than at other times during the Precambrian and, by inference, that sea level was also high at these times. It is significant that the highest peak for restored intracratonic sediment occurs at ∼1.9 Ga (Figure 2). Better resolution of these peaks must await more precise dating of the sedimentary successions.

Figure 1.

Histogram showing the cumulative aerial distribution of preserved intracratonic, passive margin, and platform sediments during the Precambrian. Data are from Table 1.

Figure 2.

Histogram showing the restored aerial distribution of intracratonic, passive margin, and platform sediments during the Precambrian. Data are from Table 1. The area was restored using a sediment half-life of 700 Myr and the exponential decay relation given by Veizer [1988].

[11] Also supporting a high sea level at ∼1.9 Ga is the widespread occurrence of submarine flood basalts of this age erupted on continental platforms. Many examples of such basalts occur in the circum-Superior orogen in Quebec, in the Birrimian in West Africa, and on the Baltic shield in Scandinavia [Arndt, 1999]. This suggests that continental shelves were extensively inundated at 1.9 Ga.

[12] A high sea level at 1.9 Ga is consistent with a superplume event at this time, as proposed by Condie [1998, 2000a]. The decrease in sea level after 1.85 Ga recorded by a decrease in preserved intracratonic sediments may reflect the formation of a supercontinent that followed and partially overlapped a superplume event [Condie, 2000b]. This is consistent with the timing of events within the superevent cycle suggested by Condie [1998].

3. Paleoclimate

[13] The chemical index of alteration (CIA), or paleoweathering index [Nesbitt and Young, 1982; Nesbitt et al., 1996], has been used to estimate the degree of chemical weathering in the source areas of shale. The CIA is calculated from the molecular proportions of oxides, where CaO is the amount in only the silicate fraction and CIA equals the [Al2O3/(Al2O3 + CaO + Na2O + K2O)] × 100 molecular ratio. The higher the CIA, the greater the degree of chemical weathering in sediment sources. For example, CIA values over 85 are characteristic of residual clays in tropical climates, and [Nesbitt and Young, 1982]. Although CIA data show considerable scatter in some stratigraphic sections, due perhaps to later remobilization of Ca, Na, or K, results indicate a major peak in CIA values of shales at ∼1.9 Ga and another at 1.7 Ga [Condie and Des Marais, 1999; Condie et al., 2000] (Figure 3). The peaks in CIA at 1.9 and 1.7 Ga suggest that paleoclimates were unusually warm at these times, a feature consistent with increased input of greenhouse gases (principally CO2) into the atmosphere. This is to be expected during superplume events.

Figure 3.

Time series of BIF, black shale, CIA in intracratonic shales, and restored intracratonic sediments in the Proterozoic compared with the distribution of global plumes from Isley and Abbott [1999]. Time series were generated by summing Gaussian distributions of unit area using mean ages and standard deviations given by Condie et al. [2000]. CIA equals the [Al2O3/(Al2O3 + CaO + Na2O + K2O) × 100 molecular ratio, with CaO representing the silicate fraction only. CIA values are given by Condie et al. [2000]. In the time series, CIA represents a sum of normal distributions whose height is determined by the error in the age of the CIA determination and the value of CIA. The relation of CIA to peak height is given approximately by CIA = 38.0 + 738 times peak height.

[14] The origin and significance of the 1.7-Ga peak in CIA is not well understood. The fact that it correlates with intracratonic sediment and black shale peaks at about the same time (Figure 3) indicates a warm climatic regime with relatively high sea level at 1.7 Ga. Isley and Abbott [1999] have proposed another superplume event at ∼1.7 Ga, and the sediment and CIA peaks seem to support the existence of such an event.

4. Banded Iron Formation (BIF)

[15] BIFs are finely laminated, chemical sediments composed chiefly of microcrystalline quartz and iron oxides. The most voluminous BIFs (Superior type) were deposited in intracratonic, passive margin, or platform basins during stands of high sea level during the late Archean and Paleoproterozoic [Simonson and Hassler, 1996]. The iron and silica in BIFs appear to have been derived from hydrothermal vents on the deep seafloor, and their deposition on shallow continental shelves and intracratonic basins requires one of two processes: (1) upwelling, which brings iron-rich waters from the largely anoxic deep basins into the oxidizing shallow water on the continents [Klein and Beukes, 1992], or (2) extensive hydrothermal plumes, depleted in oxygen and enriched in iron, associated with either or both ocean ridge systems or oceanic plateaus [Isley, 1995]. In the latter case, the hydrothermal plumes transport iron into the upper water column as they rise to a level of neutral buoyancy and spread outward because of ocean currents. In this manner the iron is carried onto continental shelves where BIF is deposited.

[16] The last major period of BIF deposition was at ∼1.9 Ga, when the large BIFs of Labrador Trough in northern Quebec, the Animikie basin in Minnesota, and the Nabberu basin in Western Australia were deposited [Klein and Beukes, 1992]. As shown by Isley and Abbott [1999], this last peak in BIF deposition correlates well with a proposed 1.9-Ga superplume event and suggests a cause and effect relationship (Figure 3). A similar correlation with a superplume event has been suggested for the voluminous BIFs at ∼2.5 Ga [Barley et al., 1997; Isley and Abbott, 1999].

[17] A superplume event can account for several features of BIF deposition. First, the enhanced submarine volcanism and hydrothermal venting associated with both ocean ridge and oceanic plateau volcanism during a superplume event may be the source of the iron and silica in BIF. Furthermore, the elevated sea level caused by a superplume event, as discussed previously, provides extensive shallow marine basins along stable continental platforms, necessary to preserve BIF against later subduction. This applies to either the upwelling or hydrothermal plume models of deposition. The end of the BIF event at 1.9 Ga may be related to either of or, more likely, a combination of the following: (1) a decrease in concentration of ferrous iron in the oceans, a feature resulting from decreasing amounts of submarine hydrothermal activity as a superplume event declines in intensity, or (2) increasing oxygenation of deep ocean waters including, perhaps, introduction of dense plumes of sulfate-rich water. After 1.8 Ga the effects of oxygenic photosynthesis, together with organic burial and a weaker hydrothermal flux, led to global sulfate deposition and to the complete disappearance of BIF. It would appear that later superplume events, including a possible event at 1.7 Ga, were insufficient to reinitiate deposition of BIF.

5. Sedimentary phosphates

[18] In many respects, what we know about the deposition of BIF also applies to many sedimentary marine phosphates. The main difference between the environments of deposition is that phosphates are deposited in biologically productive upwelling zones, whereas BIFs are not. For marine phosphate deposition we need a source of phosphorus, relatively anoxic seawater to keep the phosphorus in solution, and upwelling along continental margins (or hydrothermal plumes), which brings phosphorus-rich seawater onto continental shelves and basins that contain oxidizing water and biologic activity where phosphates can be precipitated [Cook and Shergold, 1984, 1986].

[19] Although sedimentary phosphates do not become widespread until after ∼800 Ma, some important deposits occur in the Paleoproterozoic [Cook and McElhinny, 1979]. These are in Australia in the Rum Jungle (1.9 Ga) and Broken Hill (1.9–1.8 Ga) areas, in the Animikie/Gunflint successions in Minnesota (∼1.9 Ga), in the Vayrylankyla area of Finland (2.0–1.9 Ga), and in the Yenisey province of Siberia (1.85 Ga) [Cook and McElhinny, 1979; Needham et al., 1988; Rosen et al., 1994; Nutman and Ehlers, 1998]. These phosphates were deposited at or near 1.9 Ga and hence may correlate with the proposed 1.9-Ga superplume event. The source of the increased phosphorus and widespread anoxia in seawater at this time could be from submarine hydrothermal systems associated with the superplume event. This model also explains the association of some phosphates with BIFs, in that dissolved phosphate is strongly absorbed on ferric oxides under aerobic conditions [Berner, 1999].

6. Sr Isotopes in Seawater

[20] Veizer and Compston [1976] were the first to suggest that the growth rate of continental crust can be tracked with Sr isotopes using the 87Sr/86Sr ratio of marine carbonates. Because the 87Sr/86Sr ratio of marine carbonates reflects chiefly the balance of Sr contributed from continental versus deep-sea hydrothermal sources [Veizer, 1989; Asmerom et al., 1991], relatively high 87Sr/86Sr ratios indicate sources with high Rb/Sr ratios (continental crust), whereas low ratios reflect dominant mantle input. Prior to ∼2.5 Ga, 87Sr/86Sr ratios fall near the mantle growth curve (∼0.702), indicating that the oceans were largely buffered by mantle input and that little continental crust existed, at least above sea level [Veizer and Compston, 1976]. After a major period of continental crustal growth and emergence above sea level at 2.7–2.6 Ga [Veizer and Compston, 1976; Taylor and McLennan, 1985; Condie, 1998], however, the 87Sr/86Sr ratio in seawater increased rapidly as increased amounts of continental Sr entered the oceans.

[21] In the Paleoproterozoic the 87Sr/86Sr ratio in some of the least radiogenic marine carbonates (such as those from the Albanel Formation in Quebec (1.85 Ga) and the McArthur Group in northern Australia (1.75 Ga)) increased to values near 0.706, whereas in slightly older carbonates the ratio remained relatively low (0.704) (for example, in the Coronation Supergroup in Canada (1.9 Ga)) near the mantle growth curve [Mirota and Veizer, 1994]. Perhaps the relatively low 87Sr/86Sr ratios in seawater at 1.9 Ga reflect increased mantle input of Sr from the proposed superplume event at this time, whereas the higher ratios at 1.85–1.75 Ga reflect increased input of continental Sr from a growing supercontinent [Condie, 1998]. If there was a superplume event at 1.9 Ga, why did it not decrease seawater Sr isotope ratios to even lower values? One possible reason is the large volume of continental crust formed at 2.7 Ga, which, even during a 1.9-Ga superplume event, continued to supply significant amounts of continental Sr to the oceans. Intermediate 87Sr/86Sr ratios in Mesoproterozoic marine carbonates (∼0.705) correlate with supercontinent stasis, and perhaps with minor breakup at 1.5–1.4 Ga [Condie, 2000c].

7. Stromatolites

[22] Stromatolites, layered structures thought to be deposited by microbial mat communities, are widespread in the Proterozoic with a prominent peak (or peaks) in distribution at ∼1.9–1.8 Ga. Maxima at this time occur in the number of stromatolite occurrences, in the diversity of stromatolites, and in the number of occurrences of microdigitate stromatolites [Grotzinger and Kasting, 1993; Hofmann, 1998] (Figure 4). Other investigators have not recognized this peak because data were averaged over long time intervals [Awramik, 1992; Semikhatov and Raaben, 1996]. In addition to stromatolites there are maxima in the reported occurrences of microfossils, oncoids, and chemofossils (biogenic chemical remains) at ∼1.9 Ga [Hofmann, 1998]. The peaks in abundance and diversity of these fossils at ∼1.9 Ga may reflect a combination of global warming, high sea level stands, and enhanced input of CO2 into the sedimentary cycle. All of these may be related to a superplume event at 1.9 Ga. Grotzinger and Knoll [1999] suggest that the degree of carbonate saturation in seawater may be very important in controlling stromatolite diversity, and during superplume events, seawater carbonate saturation may increase significantly. Thus a superplume event may increase both the availability of carbonate and the proportion of shallow platforms for the deposition and preservation of carbonates.

Figure 4.

Distribution of reported number of occurrences of total stromatolites and of microdigitate stromatolites during the Precambrian. Data are from Grotzinger and Kasting [1993] and Hofmann [1998].

[23] The distribution of Paleoproterozoic carbonates indicates that cement crusts, and, in particular, microdigitate stromatolites deposited in tidal flats, were a common mode of deposition of calcium carbonate during the Paleoproterozoic [Grotzinger and Kasting, 1993]. This feature appears to require Proterozoic seawater that was greatly oversaturated in CaCO3 compared to Phanerozoic seawater. The peak in reported occurrences of microdigitate stromatolites at ∼1.9 Ga (Figure 4) is particularly intriguing in that it correlates with the suggested superplume event at this time. This is consistent with enhanced CO2 input into the oceans from submarine volcanism and hydrothermal vents accompanying the superplume event, because higher CO2 means an increase in the HCO3/SO4 ratio in seawater, favoring deposition of carbonate over sulfate. Also, high sea level stands create widespread shallow tidal flats where both Ca and HCO3 ions increase in concentration in seawater because of evaporation.

8. Massive Sulfate Evaporites

[24] The first massive sulfate evaporites in the geologic record occur at 1.8–1.6 Ga, following a possible 1.9-Ga superplume event (Figure 5). The fact that widespread sulfate deposition followed a 1.9-Ga superplume event is consistent with the following sequence of events: (1) Large amounts of sulfur were injected into the oceans as sulfide during the superplume event, but only some of the sulfur was deposited as iron sulfides on the deep seafloor [Canfield, 1998]; (2) the oceans became oxic as submarine volcanic input related to the superplume event subsided; and (3) as carbonate levels decreased in seawater because of falling CO2 input from the superplumes, marine carbonate deposition became less important and Ca++ ion became available to precipitate as sulfates [Grotzinger and Kasting, 1993]. Increasing levels of sulfate in the oceans at this time are also supported by sulfur isotopes in which the range of δ34S increases (less than -20 to greater than +20) after 2.2 Ga [Canfield, 1998]. Also, the completion of a supercontinent at this time may have provided numerous partially closed basins for evaporite deposition, as it did on Pangea during the Permian and Triassic.

Figure 5.

Distribution of massive sulfate evaporites during the Precambrian. Data are from Grotzinger and Kasting [1993] and miscellaneous sources.

[25] Alternatively, the sulfate, black shale, and CIA peaks at 1.7 Ga may reflect another superplume event at this time, with all of these peaks recording widespread global warming. If so, why is there no evidence of sulfate evaporite deposition corresponding to possible superplume events at 2.7 and 1.9 Ga? There probably was not enough oxygen available to oxidize sulfur at 2.7 Ga. However, by 1.9 Ga, this should not have been a problem. If there was a 1.7-Ga superplume event, for some reason it must have caused more global warming than the 1.9-Ga event, thus leading to widespread sulfate deposition.

9. Black Shales and Carbon and Sulfur Isotopes

[26] There is a good correlation between a 1.9-Ga superplume event and the distribution of black shales, expressed either as total cumulative thickness or as the black shale/total shale ratio (Figure 3) [Condie and Des Marais, 1999; Condie et al., 2000]. The black shale and CIA peaks may reflect the combined effects of mantle superplume events and supercontinent formation, the former of which introduced massive amounts of CO2 into the atmosphere-ocean system, increasing depositional rates carbon and increasing global warming. Increased black shale deposition at these times is due to some combination of (1) increased oceanic hydrothermal fluxes (introducing nutrients), (2) anoxia driven onto continental shelves, and (3) disrupted ocean currents.

[27] The impact of a superplume upon the biogeochemical cycles of carbon and sulfur can be explored further by considering their stable isotopic records. These cycles consist of elemental reservoirs linked by processes that either transport or chemically transform these elements. Carbon and sulfur enter the surface environment by both weathering and thermal processes (i.e., volcanism, hydrothermal activity, and metamorphism). These elements are then rapidly cycled through the biosphere, which converts a fraction of their flows to reduced species by utilizing reducing power provided by weathering, thermal sources, and oxygenic photosynthesis. The amounts of organic carbon, carbonate, sulfides, and sulfate buried depend upon both the elemental fluxes through the surface environment and the burial rates in a range of sedimentary environments that favor sedimentation of either oxidized or reduced species. For example, the burial of both reduced carbon and sulfides is favored by reducing marine environments [Berner, 1983].

[28] The cycling of carbon can be monitored via an isotopic mass balance [Des Marais et al., 1992]:

equation image

where δCin is the mean isotopic composition of carbon entering the surface environment. The right side of the equation represents the weighted-average isotopic composition of carbonate (δ13Ccarb) and organic (δ13Corg) carbon being buried, and fcarb and forg represent the fractions of carbon buried in each form (fcarb = 1 forg). Over timescales of >100 Myr, δin = 5%, the average value for crustal carbon [Holser et al., 1988]. Thus, where values of sedimentary δcarb and δorg can be measured, it is possible to determine fcarb and forg for ancient carbon cycles. For example, lower values of δ13Ccarb and/or δ13Corg indicate higher values of fcarb.

[29] A similar mass balance equation applies for sulfur, as follows:

equation image

where δSin represents the isotopic composition of sulfur entering the surface environment. The right side of the equation represents the weighted-average isotopic composition of sulfate (δ34SSO4=) and sulfides (δ34SSred) being buried, and fSO4= and fSred represent the fractions of sulfur buried in each form (fSO4= = 1 − fSred). Over timescales of >100 Myr, δin = 0%, the average value for crustal sulfur [Holser et al., 1988]. Thus, for example, lower values of δ34SSred indicate higher values of fSO4=.

[30] The Paleoproterozoic interval witnessed large global excursions in δ13Ccarb values (Figure 6). The very positive δ13Ccarb values situated between 2.44 and 2.39 Ga and between 1.92 and 1.97 Ga each represent only single sedimentary basins, and therefore it is not yet established that they record global-scale phenomena [Melezhik et al., 1999; Buick et al., 1998]. However, each of the positive and negative isotopic excursions between 2.3 and 2.0 Ga are documented within multiple basins and very likely represent widespread events [Melezhik et al., 1999]. If these δ13Ccarb values reflect global excursions in δ13Cinorganic values of Paleoproterozoic seawater, then the fraction of carbon buried as organic matter (forg) varied repeatedly from less than 20% to more than 50% of the global carbon flux [Karhu and Holland, 1996]. However, at least some of the most positive δ13Ccarb values apparently developed within large hypersaline, restricted basins that also sustained abundant stromatolite growth [Melezhik et al., 1999]. During the proposed 1.9-Ga superplume event, moderately positive δ13Ccarb values indicate that forg was slightly higher than its long-term global average value, consistent with the observed abundance peak in black shales. Remarkably, after 1.92 Ga, δ13Ccarb values and forg remained relatively constant for hundreds of millions of years. This onset of stability is consistent with the idea that conditions surrounding the proposed superplume event actually stabilized the relative rates of burial of organic and carbonate carbon and thereby dampened large excursions in δ13Ccarb. For example, the sea level rise associated with a superplume event could have flooded restricted basins, coupling them more strongly to the global ocean. Increased thermal inputs of CO2 would have mitigated any local limitations in the supply of CO2 within biologically productive zones, limitations that might otherwise have led to very positive δ13Ccarb values [Melezhik et al., 1999].

Figure 6.

Carbon isotopic composition (δ13Ccarb) of Proterozoic carbonates versus age (Ga). Data are from the following sources: Each data point within the shaded area between 2.7 and 1.65 Ga represents an average of multiple analyses (see reviews by Karhu and Holland [1996] and Melezhik et al. [1999]); unmarked crosses both at 2.7 Ga and between 1.7 and 0.9 Ga are 100-Ma running averages of data compiled by Des Marais [1997]; section between 1.7 and 1.5 Ga is from Brasier and Lindsay [1998]; data labeled “H” and “B” are from Hall and Veizer [1996] and Buick et al. [1995]], respectively; the shaded field of data between 0.5 and 0.85 Ga is from a review by Kaufman and Knoll [1995]. Question marks highlight data points that represent single basins only and therefore might reflect only regional, rather than global, isotopic excursions [see Melezhik et al., 1999]. The oval encloses data representing the period surrounding the proposed 1.9-Ga superplume event.

[31] The range of δSred values increased between 2.5 and 2.2 Ga and is consistent with an increase in seawater sulfate levels [Knoll and Canfield, 1998] (Figure 7). Because bacterial sulfate reduction utilizes 32S preferentially over 34S [Harrison and Thode, 1958], sedimentary sulfides typically have lower δ34S values than their coeval sulfates. Isotopic mass balance considerations require that the weighted average of the δSO4= and δSred values of sulfur species being buried must equal δSin, which is typically 0 (see above). Between 1.9 and 1.8 Ga (circled data, Figure 7), δSred values become more negative, indicating both that a substantial sulfate reservoir existed and that the burial rate of sulfate increased relative to that of sulfide during that time interval. This isotopic trend is consistent with the observed peak in the distribution of sulfate evaporites between 1.8 and 1.6 Ga (Figure 5).

Figure 7.

Sulfur isotopic composition (δ34SSred) of Proterozoic sedimentary sulfides versus age. The oval encloses data representing the time period of a 1.9-Ga superplume event and its aftermath. Adapted from Knoll and Canfield [1998].

[32] The carbon and sulfur isotope trends in Figure 6 and 7 are consistent with a high sea level stand at 1.9 Ga. This is followed by a sea level decline and a decline in platform sedimentation in favor of shallow-to-emergent coastal environments that accumulated more oxidized, evaporitic (sulfate) sedimentation. This view is corroborated by the decline, from 1.9 to 1.7 Ga, in stromatolite diversity and in the abundance of preserved banded iron formation and black shale and by declining δSred values. The trend toward greater rates of sulfate deposition between 1.8 and 1.7 Ga is indicated both by increased abundance of platform sulfate-rich evaporites and generally lower δSred values.

10. Cross-Correlation Analysis

[33] During superplume events at 480, 280, and 100 Ma, sea level rose and platform-type sedimentation became widespread [Larson, 1991b]. Thus, in the Precambrian we should also expect continental inundation to accompany superplume events. However, the view that Precambrian continents were flooded during superplume events presupposes either that plate motion rates increased greatly during Precambrian superplume events or that large volumes of thick oceanic plateaus were formed within ocean basins. Both types of events are difficult to document directly, either from paleomagnetic data or from geochemical studies. Furthermore, the elevation of continents during supercontinent formation might produce a reduction in continental surface area that would counteract plume-related decreases in the volume of the ocean basins. Therefore a cross-correlation analysis of time series of black shales, CIA, and intracratonic sediments is a first-order test of the idea that Precambrian superplumes either accelerated plate motions and/or created large volumes of thick oceanic plateaus.

[34] Condie et al. [2000] report a high correlation between the abundance of black shale and a high chemical index of alteration (CIA) in the Precambrian record. Cross-correlation analyses of abundance of black shale and intracratonic sediment have correlation coefficients with confidence levels between 87 and 99.9% (Table 2). Six out of nine of the cross correlations have confidence levels of 98% or higher, with best fitting time differences of 42 Myr or less. These time differences are the number of years subtracted from (or added to) the ages in the first time series in order to produce the highest possible correlation between the first and second time series. Because the mean error of the sediment ages is 62 Myr, these time differences are effectively zero. Thus the deposition of black shale and intracratonic sediment is highly correlated during the Proterozoic. The cross correlations between the time series of CIA and intracratonic sediment abundance have confidence levels of 98–99.9%. Two out of three of the cross correlations have best fit time differences of 39 Myr, which again is less than the mean error (62 Myr) and, hence, effectively zero time difference. Thus high levels of atmospheric CO2 and deposition of platform sediments are highly correlated during the Proterozoic. Overall, these results suggest that Proterozoic superplume events were associated with either accelerations in plate motion and/or formation of large volumes of oceanic plateaus.

Table 2. Results of Cross-Correlation Analysis of Time Series of Black Shales, Intracratonic Sediments, and CIA for Rocks of >600 Maa
Time SeriesDelay, MyrCorrelation CoefficentConfidence Level, %
  • a

    Black shales are weighted according to the total thickness of the black shale (black shale weight (wt) 1) and the overall proportion of black shale in a shale sequence (black shale weight 2). Intracratonic sediments are weighted according the surface area of the sediments (sediment weight 1) and the restored surface area of sediments (sediment weight 2). The primary peak heights in the time series are related to the errors in the ages of black shale and intracratonic sediment. Thus well-dated sequences have higher peak heights than poorly dated sequences. The significance level of each correlation coefficient is calculated from 1000 simulations with random numbers of the correlation of each time series with a randomly generated time series with the same spectral characteristics [Isley and Abbott, 1998]. Time differences are derived from shifting one time series relative to the other in time until the highest possible correlation between the two time series is achieved.

CIA versus sediment310.7698.0
CIA versus sediment wt 1390.7998.0
CIA versus sediment wt 210990.8499.9
Black shale with sediment60.8999.9
Black shale with sediment wt 150.8299.9
Black shale with sediment wt 2160.7999.8
Black shale wt 1 with sediment20.7598.0
Black shale wt 1 with sediment wt 1420.7196.0
Black shale wt 1 with sediment wt 212800.6482.0
Black shale wt 2 with sediment20.7999.7
Black shale wt 2 with sediment wt 11250.7699.1
Black shale wt 2 with sediment wt 212750.6787.0

11. Discussion and Conclusions

[35] Although the results of this study support the existence of a superplume event at 1.9 Ga, they do not prove the existence of such an event. Consistent with a superplume event at this time are the following: (1) high sea level as inferred from the relative abundance of intracratonic, passive margin, and platform sediments; (2) a relatively high abundance of black shale; (3) a peak in CIA (chemical index of alteration) in shale implying unusually warm paleoclimates; (4) a peak in the abundance of BIF; (5) a peak in the abundance of shallow marine phosphate deposits; and (6) a peak in the number of reported occurrences and in the diversity of stromatolites. Although any one of these observations may be explained by different processes, it is the coincidence of all of them at 1.9 Ga that is the most compelling evidence for a superplume event [Condie et al., 2000].

[36] The peak in distribution of shallow marine sediments at 1.9 Ga suggests a corresponding high in sea level. If this reflects a superplume event, a drop in sea level during the interval of ∼1.85–1.7 Ga may be due to supercontinent formation, which followed the superplume event [Condie, 2000b]. Increased black shale deposition and preservation at 1.9 Ga is due primarily to anoxia driven onto stable continental shelves. The black shale and CIA peaks at 1.9 Ga may reflect introduction of massive amounts of CO2 into the atmosphere-ocean system, increasing depositional rates of carbon and increasing global warming. Such a massive injection of CO2 into the atmosphere is again consistent with a superplume event [Condie et al., 2000]. The absence of major carbon isotope anomalies in seawater at 1.9 Ga indicates that even though absolute burial rates of both reduced and oxidized carbon were accelerated at this time, their relative burial rates remained similar to those observed throughout most of the geologic record, including today [Condie et al., 2000].

[37] Increased BIF and marine phosphate deposition at 1.9 Ga also may be due to a superplume event. Increased rates of deposition of these sediments may be due to increased input of iron and phosphorus into deep anoxic oceans by submarine hydrothermal activity related to a superplume event, followed by upwelling (or spreading of hydrothermal plumes) into shallow, oxidizing seas on continental shelves. The slightly low 87Sr/86Sr isotope ratios in seawater at 1.9 Ga may reflect increased mantle input of Sr from a superplume event, whereas higher ratios at 1.85–1.75 Ga reflect increased input of continental Sr from a growing supercontinent. The fact that a superplume event at 1.9 Ga did not significantly decrease seawater Sr isotope ratios may be due to the large volume of continental crust formed at 2.7 Ga, which continued to supply significant amounts of continental Sr to the oceans, thus buffering the effect of mantle Sr related to the superplume event. A peak in reported occurrences and diversity of stromatolites in general and microdigitate stromatolites at ∼1.9 Ga requires seawater on continental shelves that was greatly oversaturated in CaCO3 and warm climates. Both of these observations are consistent with high sea level stands and enhanced CO2 input into the oceans from submarine volcanism and hydrothermal vents accompanying a 1.9-Ga superplume event. The first massive sulfate evaporites in the geologic record at 1.8–1.6 Ga reflect oxidizing conditions and greater availability of Ca++ ions as carbonate deposition declined during superplume waning.


[38] The authors greatly appreciate the help of Dustin Smyth in compiling and synthesizing the sediment data used in estimating Proterozoic sea levels. We also acknowledge the NASA Exobiology Program and the NASA Astrobiology Institute. This is Lamont-Doherty Earth Observatory contribution 6114. The paper was substantially improved by in-depth reviews by Hubert Staudigel and two anonymous reviewers.