The silica cycle in the Precambrian

0016-7037/92/$5.00

Geuchimica er Cosmochimica Acla Vol. 56. pp. 3265-3212 Copyright 6 1992 Pergamon Press Ltd. Printed inU.S.A.

+ .oO

The silica cycle in the Precambrian RAYM~NDSIEVER

Department of Earth and Planetary Sciences, Harvard University, Cambridge, MA 02 138, USA (Received March 19, 1991; accepted in revisedform August 23, 1991) Abslract-Whereas

the modem silica cycle is heavily dependent on the biology of silica-secreting organisms such as the diatoms, the Precambrian silica cycle was dominated by inorganic reactions among dissolved silica, clay and zeolite minerals, and organic matter. There is no evidence for deposition of a layered amorphous silica but abundant evidence for diagenetic silicification in the Neoproterozoic, the time period chosen for this analysis. The influx of silica to the Neoproterozoic ocean would have been governed, as today, by the balance among tectonism, weathering, and hydrothermal input. The evidence of Neoproterozoic rocks favors tectonic and weathering regimes not very dissimilar to those of the early Phanerozoic and hydrothermal inputs significantly altered at certain periods. One flux would have been very different: the present diffisional influx of silica From interstitial waters into the oceans would have been altered to an efflux from the oceans to interstitial waters. Reactions of dissolved silica with inorganic phases would have controlled silica concentrations at a level of about 60 ppm. This figure is based on experimental determinations of solubility and silica sorption on clay minerals and zeolites. As silica in interstitial waters was lowered by opal-CT precipitation, a diffusional efflux from the ocean into the sediment was set up. Silicification would have started in early diagenesis as opal-CT precipitated from interstitial waters. Diffusion of silica from overlying seawater declined as the sediment was buried a few hundred meters and the diffusion path lengthened. During this still relatively shallow burial, silicification continued, but this time powered by hydrologic transport of dissolved silica from coastal plain sediments during transgressions and regressions. Thus the major removal of silica from the Neoproterozoic ocean took place by diagenetic reactions. THE SILICA CYCLE IN THE PRECAMBRIAN

mately an order of magnitude lower than the saturation value of quartz. The siliceous sediments deposited in the ocean initially as biogenic silica are transformed by diagenetic processes into bedded and nodular cherts and siliceous mudstones (see, e.g., CALVERT,1974; RIECH and VON RAD, 1979; MALIVAand SIEVER, 1989). The effect of macroevolution of silica-secreting organisms during the Phanerozoic was profound, influencing both the nature of the biogenic material and the environments of sedimentation ( MALIVAet al., 1989). During the Cenozoic, diatoms controlled silica production but at earlier times, before the diatoms had evolved, radiolaria and sponges were the dominant organisms (&EVER, 199 1) . Seawater concentrations and sediment types varied accordingly. The geochemical cycle of silica in the Precambrian was different from that of the Phanerozoic in one main respect: there was no biological precipitation of silica. The silica-secreting plants and animals of the Phanerozoic: radiolaria, sponges, diatoms, and vascular plants had not yet evolved. There is no firm evidence that any Precambrian algae or bacteria precipitated silica. [ALLISON and HILGERT ( 1986) have reported algal silica remains from the Cambrian, which based on more recent evidence is dated as Riphean ( KAUFMAN et al., i 992 ) , but there is some doubt that these materials were originally silica.] Likewise, there is no evidence that Ediacaran or other metazooan forerunners secreted silica. Thus, for this early time in geologic history, we must reconstruct a cycle primarily from inorganic reactions, although considering interactions with the biosphere insofar as they result in silica-organic matter reactions. For the bulk of this paper I discuss mainly the Neoproterozoic, the interval from 1,000 to 540 Ma. In the conclusion

THIRTY-THREEyears ago I wrote my first paper on the silica cycle. I was starting a year’s visit to Harvard, which I chose for only one reason: Bob Garrels had just gone there the year before. I had known Bob for some years from his earlier days at Northwestern, primarily through the Illinois-Iowa-Wisconsin Tri-State field trips, which were a must for every geologist of that time and place. But only when I began to work with him day by day did I start to understand the depths of his insight, the breadth of his interests, and, not least important, the sheer fun of working with him. It is fitting to contribute to this symposium a paper about the Precambrian, a time in Earth history that was a lifelong interest of Bobs. It is also appropriate that it concerns geochemical cycles, for it was during the early years of our association that I saw his interest in cycles deepen and move him into the studies in which he made some of his greatest contributions. SECULAR CHANGE IN THE SILICA CYCLE The modem silica cycle is fairly well understood and has been summarized by WOLLASTand MACKENZIE ( 1983). The influx to the oceans of silica from rivers, hydrothermal sources, and diffusion from sediment interstitial waters is largely balanced by the efflux of biochemically precipitated silica, primarily by the diatoms (Fig. 1). Becauseof the efficiency of the diatoms in precipitating silica, seawater concentrations are very low, ranging from less than one part per million (ppm) in most surface waters to regions of limited extent where bottom waters may reach lo- 15 ppm. The bulk of the oceans, at about 1 ppm, is poised at values approxi3265

3266

R. Siever COMNMAL WEATHERING WXCANISM

\

PASSlVE MARGIN

0

MD-OCEAN RIDGE

slmoucrlON ZOrE

FIG. 1. The present-day silica cycle (Modified from WOLLAST and MACKENZIE, 1983). Fluxes are given as 10” g Si/yr.

I comment on the Archean. I start with a resume of the occurrence of siliceous sediments in the later Proterozoic. NEOPROTEROZOIC

SILICEOUS

SEDIMENTARY

ROCKS

There is virtually no evidence of any extensive primary, inorganic precipitation of chert in later Proterozoic sedimentary rocks. (There are some quartz crystals of evaporite origin in the Bitter Springs lacustrine beds; SOUTHGATE,1986.) With the possible exception of some banded iron formation, field observations and microscopy indicate diagenetic replacement chert to be the rule. Replacement cherts display relic inclusions of carbonate and replaced detrital grains as well as silicified cross-bedding and other sedimentary structures. We can imagine the properties of a supposed amorphous silica sediment precipitated directly from a seawater with silica at a concentration of about 120- 140 ppm silica. Such a sediment would show primary bedding, would be in elastic-starved basins, and would look like the finely laminated silica in Phanerozoic bedded cherts, except for the lack of any silica-secreting fossils. We find no such rocks. Carbonate rocks were the chief loci of silicification; the carbonate environments in which silicification commonly occurred were dominantly peritidal to shallow subtidal (KNOLL, 1982, 1985; MALIVAet al., 1989). Few deeper water basinal limestones are known and the few that exist are not silicified. Shallow-water quartz arenites are commonly cemented with secondary quartz but do not show the kind of nodular or replacement characteristics of cherty silicification of carbonates. GEOLOGICAL

CONTROLS

ON THE SILICA

CYCLE

The primary control on the influxes of silica to the oceans is tectonic. To evaluate the size of Precambrian silica inputs we can compare the state of the continents and oceans then with their state now. Although rates of Proterozoic tectonism and volcanism are controversial, there is no firm evidence that, on the average, such rates were significantly higher in the Neoproterozoic. The range of sandstone types reflects about the same range of tectonic states as the Phanerozoic (with the possible exception of more abundant quartz arenites that could represent extraordinary weathering and/or reworking rates). In addition, rates of sedimentation and subsidence in Neoproterozoic basins seem compatible with those of later times, although GROTZINCER and ROYDEN ( 1990) have suggested that earlier Proterozoic elastic thicknesses on

the Slave craton were 12 f 4 km, corresponding to high thermal gradients and, ultimately, higher rates of subsidence. Thus I make the assumption for the Neoproterozoic that the ranges of globally averaged tectonic states and rates of changes of continents and oceans basins, on the average, were much like those of the following early Phanerozoic. But average states do not tell the whole story. Periodic perturbations of tectonic regimes may have imposed large departures from the average. For example, ASMEROMet al. ( 199 1)) on the basis of isotopic data, believe that hydrothermal influxes peaked at 830 Ma, at about the time of the Pan African and related orogenies. They also correlate high values of continental erosion rates, at about 570 Ma, with a time of large-scale continental collisional orogenies. Such perturbations in global tectonic activity would entrain changes in the kinds and global distribution of weathering, transport, and sedimentation regimes. River influx of silica to the oceans, dependent on weathering and erosion rates, will reflect higher weathering rates in humid equatorial belts and in elevated land masses associated with major orogenies. Increased volcanism at active plate convergence boundaries during times of rapid plate motions also contributes to greater weathering rates, for mahc volcanics weather rapidly. During times without polar icecaps, weathering would be heightened. Notwithstanding higher sea levels during ice-free times, which would flood continental lowlands, all slightly elevated areas of the continents would be exposed to weathering and erosion. During major glaciations, although there was much mechanical erosion by glaciers, rates of chemical weathering would be minimized because of the large continental areas covered by ice. These comparisons can be made meaningful, however, only by evaluation of the proportion of continents at low and high latitudes. Continental drift reconstructions place the largest land areas of the later Proterozoic near the equator; this has significant implications for river influxes. Assuming latitudinal climate belts something like today’s, much of the land area of the globe would lie in latitudes of heaviest precipitation, which would maximize weathering, erosion, and river transport. The complementary desert regimes might also have been extensive but would not have strongly influenced silica contributions of rivers. Hydrothermal influx of silica to the oceans, dependent on plate spreading rates, may have been considerably greater in certain times in the Neoproterozoic than in the early Pha-

Geochemistry of silica in Precambrian time nerozoic ( ASMEROM et al., 199 1). Yet Precambrian polar wandering curves, with all their uncertainties, seem to reflect a “normal” spreading regime. At best, average rates of spreading remain uncertain. Nevertheless, within the range of variation, the general pattern and order of magnitude of all ofthe influxes and effluxes would have been characteristic of much of the Proterozoic and significantly different from the Phanerozoic. Much of that difference can be explained in terms of organic macroevolution.

provided by low oxygen levels would have been more pronounced earlier in the Proterozoic. More isotopic data and analysis of the Proterozoic like that of KNOLL et al. ( 1986) will help us learn the history of organic matter. Regardless of the details, Proterozoic sediments contained a range of organic matter concentrations not unlike those of the Phanerozoic, and, as I will show, organic matter is implicated in the silicification process. SUMMARY OF PROTERO~I~ INF’LUXRS TO THE OCEANS

LATER PROTEROZOIC ORGANISMS Although we have no evidence of silica-secreting organisms, organisms were abundant long before the later Proterozoic. Eukaryotes had evolved by 1,800- 1,700 Ma or earlier, and in the Vendian, the Ediacaran fauna appeared. Cyanobacteria and other bacteria were abundant and strongly affected carbonate sedimentation in the oceans, with the accumulation of stromatolites and algal mats. It is likely that land surfaces were already colonized by many different kinds of bacteria and fungi. The evidence of Proterozoic coal beds indicates that large biomasses accumulated locally, perhaps in the same kinds of hydrologic environments as today. Because of the possible importance of silica-organic matter interactions, it is relevant to try to assess the kinds and amounts of organic matter in sedimentary environments. It seems likely that eukaryotic species followed general evolutionary patterns of later times: first, origin of species; second, species radiation and exploration of different environments; third, large growths of biomass in particular environment. By the later Proterozoic, as in the Phanerozoic, there would have been an abundant global biomass as well as large local biomasses in some environments. Dissolved organic carbon would long since have made its appearance in rivers and the ocean. An increasing biomass would likely have been echoed by increasing organic carbon burial, assuming the partition between burial and oxidative degradation remained more or less constant and that there was consequent atmospheric oxygen production. Because of the feedbacks in the carbon system, an oxygen pressure in the atmosphere somewhat lower than that of today would have contributed to less oxidative d~om~sition and thus greater pre~~ation and burial. As oxygen in the atmosphere slowly rose during the Proterozoic, the ratio of degradation to burial would have increased. Because it seems likely that the middle and later Proterozoic were times when the pressure of oxygen in the atmosphere was increasing, the inhibition of organic matter degradation

A quantitative estimate of the silica influxes to the ocean in the later Proterozoic (Fig. 2) depends on comparisons with the present continental regime, which may not be representative of much of the Phanerozoic. Were later Proterozoic average rates of river runoff and continental weathering the same as today’s? Perhaps they were more comparable to those periods of the Phanerozoic when most continents were near the equator and runoff high. Comparing also with respect to tectonic state, we can ask how fast were average rates of weathering and runoff at around 600-700 Ma, when large fractions of the continental surface were elevated above sea level and dominated by mountain-balding. Were they more similar to those of the Permo-Triassic, when large areas of the continents were both high and close to the equator, than they were to those of today, when continents are high both in elevation and latitude? Weathering The evidence from sedimentary rocks on the intensity of weathering is neither clear nor quantitative. As noted above, many thick quartz arenites of this period are suggestive of extensive weathering. On the other hand, BASU ( 198 1) has presented some data that before K-chelating land plants developed in the early Phanerozoic, K-feldspar was more resistant to weathering than plagioclase, thus lowering the average rate of production of free silica by silicate weathering. This and other arguments on higher plants and weathering would lead us to expect large changes in weathering, not at the Precambrian-Cambrian boundary, but at the time of extensive vascular plant colonization of the land surface in the late Silurian and Devonian. Oxygen levels in the later Proterozoic atmosphere lower than at present (HOLLAND and BEUKES, 1990) would have implied lower rates of oxidative weathering than at present but this might not have affected silica fluxes. Experimental

CONTINENTAL WEATHERING DIAGENETIC SDfWTlON AND REt-CTION---; \

-WSIM FIG.

MARGIN

3267

MID-OCEAN RIDGE SUB0

2.The later Proterozoic silica cycle for the time interval 570-600 MaBP.

3268

R. Siever

evidence suggests that silicate dissolution may be unaffected or even enhanced in a low-oxygen atmosphere, primarily because ferrous iron is relatively soluble ( SIEVERand WOODFORD, 1979). The evidence from the rocks suggests that weathering in the later Proterozoic was not much different in kind or intensity from that of various times in the Phanerozoic, the global influx from rivers varying with continental elevation and latitude. I show two possible values for river influx: ( 1) 250 (IO’* g Si/yr) for the period around 570-600 MaBP, when continental erosion was high as a result of collisional orogeny (Fig. 2 ) ; and ( 2 ) 300 for the period 8 1O-840 MaBP, at the time of high hydrothermal influx during parts of the Pan-African orogeny (Fig. 3). Both values are larger than the present influx of 203; these high values are primarily attributable to the high proportion of the continental surface near the equator. Hydrothermal and Diffusional Influxes

Influxes of silica to the Ocean from sources other than continental weathering may have been different from those of today. As noted earlier, the hydrothermal influx was much greater in the period 8 lo-840 MaBP. In Figs. 2 and 3 I show estimates of hydrothermal fluxes that are very high, 350, for the period 810-840 MaBP, and, for the period 570-600 MaBP, 100, somewhat elevated over that of the present. The hydrothermal contribution may have been slightly higher at the latter time because of a heat flow slightly elevated over the present. A much more significant difference lies in the diffusional influx to the ocean from sea floor sediments. At present, silica diffuses into seawater because the concentration of silica in interstitial waters, derived from mineral, diatom, and radiolarian dissolution, is much higher than the overlying seawater. In the later Proterozoic there would have been no sedimentary biogenic silica to dissolve in interstitial waters. In addition, seawater was much higher in silica concentration than today because of the absence of silica secreters. As a consequence, silica would have diffused into, not out of, the sediment. The diffusional flux from the sediment at present is driven by large concentration differences in the range 60100 ppm in many regions. The diffusional flux into the sediment in the Proterozoic would likely have been smaller, estimated here at about 50 ( 10 ‘*g/yr). This value is based on estimates of sea and interstitial water silica concentrations that are in turn dependent on reactions of dissolved silica with various solid phases as will be discussed.

REACHONS GOVERNING OCEANIC SILICA CONCENTRATIONS In the absence of biological silica-secreters, the phases and reactions governing seawater silica concentrations were those of dissolved silica with clay minerals, zeolites, silica phases, and organic matter. Sorption of silica on different kaolinites, illites, chlorites, and smectites has been shown to depend on mineralogy, surface area, pH, and cation concentrations ( SIEVER and WOODFORD, 1973 ). Additional experimental data on zeolites, using the same methods, are shown in Table 1 and Fig. 4. Taken together and considering the relative abundance of the different clay minerals and zeolites in the ocean, these data suggest a kinetic sorption quasi-equilibrium of about 60 ppm (approximately 1 mM). Silica-organic matter reactions are not considered in this global evaluation, although they may be locally important. The total amount of dissolved or particulate organic matter in the oceans would have been dwarfed by the silicate phases and would not have significantly affected global silica concentrations. Another important reaction that would have participated in the control of seawater silica concentration is the crystallization of opal-CT. This phase, through which most chert passes on its diagenetic way from opal-A to quartz ( MALIVA and SIEVER, 1988), precipitates at a lower value than 60 ppm. Experiments and analysis of natural occurrences indicate that at 25’C the upper bound for saturation with respect to disordered opal-CT is around 60 ppm and the lower bound for well-ordered opal-CT is around 25 ppm ( KASTNER et al., 1977; KASTNER and SIEVER, 1983). In sea and interstitial waters at these concentrations, opal-CT nucleates and grows as a kinetically favored species over chalcedony and quartz, the stabler phases. In very slow sedimentation rate environments we may expect nucleation at the sedimentwater interface. In environments with rapid sedimentation rates, the opal-CT would start to grow in early diagenetic interstitial waters. Either way, the effect is to draw down seawater silica concentrations, the first by direct precipitation and the second by diffusion into the sediment. Opal-CT grows slowly at the low temperatures of the deep sea floor, even in interstitial waters with high silica concentrations found in today’s siliceous oozes. Even in warm, shallow seas, most opal-CT in the later Proterozoic would have crystallized in interstitial waters. Complexing with and sorption by organic matter may have been an important reaction in Proterozoic seawaters. We know from the sedimentary record that abundant refractory and labile solid, fluid, and dissolved organic compounds were

CONTINENTAL WEATHERING \

DWjENETlC SORPTION AND REACTION - -1

M5SIVE FIG. 3.

vouXuI.sM

MARGIN

The later Proterozoic silica cycle for the time interval 8 IO-840 MaBP.

Geochemistry of silica in Precambriantime Table

1.

Silica

sorption

Sample

of

clay

and zeolite

Solubility

Sorption

2302

Clinoptilolite

35.5*

50*

2280

Clinoptilolite

61

70

2269

Clinoptilolite

36

39

2273

Clinoptilolite

55

69

2283

Clinoptilolite

51

57

2298

Phillipsite

39

47.5

2299

Phillipsite

48

69.5

2300

Phillipsite

25

33

2289

Erionite

55

70

2291

Erionite

28

44.5

2295

Mordenite

19

23

2297

Mordenite

34.5

46.5

2263

Analcime

18

29.5

Samples given

levels

run

for one year in O.OlM NaNC03 buffered

in Siever *Sample Samwle

and Woodford

3269

minerals

equilibrium

solution

by methods

(1973).

run for one month locations:

2302

Ft.

2280

Malheur,

2269

Campbell,

2273

Barstow,

2283

Mt.

2298

Kirkland,

Arizona

2299

Shoshone,

California

2300

Pine Valley,

2289

Needle

2291

Malheur,

2295

Pismo Beach, California

2297

Trinity

2263

Barstow,

Ward,

Washington Oregon Alabama California

Green,

Utah

Nevada

Peak, Nevada Oregon

Basin, Nevada California

present in sedimentary environments. We have known for a long time of such compounds as alkyl henzincs, naphthalene, monoterpenoids, methyl esters, and aromatic hydrocarbons

that are extractable from Precambrian

sediments. Rapidly

growing knowledge of later Proterozoic organisms suggests a large range of organic compounds, including several occur-

rences of hydrocarbons, steranes, triterpanes, and extended acyclic isoprenoid alkanes (SUMMONSand WALTER, 1990 1.

3270

R. Siever The extent of silica sorption and opal-CT precipitation would have depended on silica concentrations. In the absence of biological forcing of silica, the silica concentrations of the bulk of the oceans would have been fairly uniform, poised at the 60 ppm level. The silica concentrations of nearshore parts of the oceans might have varied, especially in evaporite environments, where concentrations might have reached opal-A saturation. Early Diagenesis

50-

25-

1.

d_m

-SMECTtTE

= %ZWE - CHALCEDOW -QUARTZ FIG. 4. Silica sorption levels of clay and zeolite minerals.

It is likely that the amount of organic matter in a continental margin marine sediment would be proportional to the sedimentation rate of the tenigenous sediment, an approximation that has been made for the Phanerozoic by BERNER( 1990). Sorption experiments in this and other laboratories point to efficient sorption of silica on natural particulate organic matter as well as a number of refractory solid organic compounds. Complexation experiments confirm indications from biochemical transport of silica that there are silica-organic complexes with such compounds as tartrate and catechol that have fairly strong bonding energies (BENNETT, 199 1; BENNETT and SIEGEL, 1987). In addition, innumerable experiments with silica-based chromatography have shown that opaline silica phases sorb dissolved organic carbon species. Finally, there are biochemical reactions with silica that suggest that some bacteria promote silicification ( BIRNBAUM and WIREMAN, 1984, 1985). Organic degradation products also promote silicification. SILICA EFFLUX

FROM THE OCEAN

The foregoing discussion implies that the silica efflux from the oceans in the absence of biological silica production must have been almost entirely diagenetic. The bulk of the silica would have been sorbed by clay minerals and zeolites, which were widely distributed over the sea floor, then as now, In the rapid sedimentation environments of continental margins the sorption on these phases would have taken place during early diagenesis. Over the majority of the deep sea floor, with slow sedimentation rates, silica sorption may have taken place at the sediment-water interface. The same considerations apply to opal-CT precipitation, although there are processes, discussed later, that might enhance opal-CT precipitation in nearshore, shallow environments.

A sequence of early diagenetic reactions would have been responsible for the removal of silica from the oceans. For these reactions I assume a geological environment stable for at least IO5 years with no significant changes in local sea level. In the first of these reactions, dissolved silica would sorb onto clay minerals and organic matter and form an amorphous silica surface phase, both in the water column and at the sediment-water interface. After sediment burial, microbial activity would degrade more labile organic matter to carbon dioxide and/or methane. As organic matter decreased or disappeared, some of the silica sorbed on that organic matter would be transformed from a sorbed layer state to a more stable polymer by hydrogen bonding to (OH) groups, as shown for wood silicification by LEO and BARGHOORN( 1976). A further transformation would produce crystallites of opal-A, leading to well-preserved organic microfossils like those discovered by BARGHOORNand TYLER ( 1965). The most labile fractions of the organic matter would be freed to solution as the organic matter degrades. In the next step, opal-CT would nucleate in pore spaces and draw down dissolved silica concentrations in interstitial waters. As a result of lower silica concentrations, clay minerals would desorb silica. This desorption may have been enhanced by the effect of lowered pH values, which would have resulted from the increased amount of carbon dioxide liberated by organic matter degradation. As silica was depleted from interstitial waters to levels well below 60 ppm, diffusion downward from seawater into the sediment replenished the silica that was withdrawn by opal-CT precipitation. Nodule growth and silicification in the shallowly buried sediment would have progressed rapidly while supported by desorption from silicates and degrading organic matter, but much more slowly as silica concentrations were supported mainly by downward diffusion from seawater. The rate would further decrease as the sediment subsided and the diffusion path lengthened. Much of the silicification would have been localized in carbonate rocks for several reasons. It has been shown that silica replacement of carbonate is particularly efficient because of the differential pressure solution and force of crystallization exerted by silica phases growing at the interface with carbonate minerals (MALIVAand SIEVER,1988). In addition, abundant organic matter is associated with carbonate shallow-shelf sediments of this age (KNOLL, 1985). Finally, the kinetics of the opal-A to opal-CT transformation are fastest in carbonate sediments and slowest in clay mineral sediments (KASTNER et al., 1977).

Geochemistry of silica in Precambrian time Early-middle Diagenesis Additional silicification would have proceeded during early-middle diagenesis, which 1 define here as the period during which diffusional contact with seawater has been severed. The lack of effective diffusion results from the thickness of sediment burial, in the range of 500 m, but well before deeper burial to several kilometers. During early-middle diagenesis, the mobility of sea level, accompanied by onlapping, offlapping, and progradation, which produce transgressions and regressions, results in hydrologic movements, chemical transport, and diagenesis of silica ( SIEVER, 1983 ). In the mixing zone fresh and seawater carbonate may become undersaturated and silica supersaturated in interstitial waters ( KNAUTH, 1979 ) . Mixing is complex because of hydrostatic heads from land, differential permeabilities, and Kohout convection from subsea into subland waters ( SIMMS, 1984). Because of transgressions and regressions, the mixing zone oscillates with time. During regressions, early diagenetically precipitated silica and authigenic and detrital silicates dissolve in aggressive meteoric waters moving downdip in shoreline deposits. Clay minerals release silica in the relatively low ionic-strength, low-pH environment of coastal plain subsurface waters. Silicasaturated or supersaturated meteoric waters are slow to nucleate opal-CT on clays or sands because of kinetic hindrance. But when the waters reach and mix with interstitial waters of a nearshore carbonate facies, the kinetics change and the precipitation of opal-CT speeds up. The overall process is the near-shoreline removal of silica from silicates and highly disordered surface silica phases and precipitation in carbonates. Evaporites may play a role in diagenetic silica precipitation or even direct primary precipitation. At a 10: 1 water volume reduction, dissolved silica would reach about 600 ppm (0.1 M), about five-fold supersaturated with respect to amorphous silica. Given these high silica values, silica would be both precipitated as opal-CT in the evaporite carbonate belt (landward of halite-gypsum) and sorbed on silicate silts and clays on shoreline detritals coeval with evaporites. According to the above, most control over the silica cycle is exerted by diagenesis along shallow, passive continental margins. Deeper basinal turbidites show little silicification. Sedimentation rates of these deposits were fast, allowing relatively short early diagenesis times for silica sorption. The mineral composition would have inhibited precipitation of opal-CT. Sorption on deep-sea pelagic clays would have played some role in the silica cycle. Because of the slow rates of sedimentation of pelagic clays, the clay mineral and zeolite surfaces would have been saturated by sorbed silica before burial. Yet there is indirect evidence that deep-sea diagenesis did not account for much silica efflux. The little that we know of later Precambrian deep-sea deposits comes from ophiolites; those of the Proterozoic are not known to contain bedded siliceous sediments (P. Hoffman, pers. comm.) . If this is so, it would indicate little or no deep-sea silica precipitation. One of the implications of this kind of silica cycle is a relationship to atmospheric oxygen evolution. Silica-organic

3271

matter reactions seem to have played an important part in diagenetic silica precipitation. The amount of organic matter in the sediment is a balance between production and degradation. In a great many later Proterozoic sediments, average or low organic matter content probably reflects dominance of bacterial degradation. The greater amount of organic matter in some nearshore sediments would imply predominance of production over degradation. This predominance presumably would be favored by locally relatively low atmospheric oxygen levels, which would inhibit bacterial oxidative decay, assuming that abiogenic oxidation is negligible. In these environments, silica precipitation would be favored. But if globally an appreciable amount of organic matter is buried, the oxygen pressure of the atmosphere would be increased, other things being equal. At the same time, the relative biomass of oxidative bacteria would be increased as oxygen levels rose. In this complicated set of feedbacks, the carbon isotopic composition is affected, as shown by KNOLL et al. ( 1986), who suggested that a major carbon isotopic excursion during the Riphean-Vendian correlates with a period of globally increased organic matter burial. This would also correlate with more extensive silicification. In this odd way, diagenetic silicification and the silica cycle are tied to the carbon cycle. The Archean The Archean presents different problems than the later Proterozoic. The earlier Proterozoic can best be seen as a long transition period. The intensity and distribution of tectonism and volcanism in the Archean, as well as climate, all of which would have affected continental weathering rates, may have differed, although field evidence continues to show that surface processes were much the same in kind as today’s (PETTIJOHN, 1943; HOFFMAN, 1980; GROTZINGER, 1986). It may be that volcaniclastics and hydrothermal processes dominated the cycle then. We also have to consider the stages of bacterial evolution and the oxygen and carbon dioxide pressures in the atmosphere. The only organisms were prokaryotes and perhaps non-fossilized forms ancestral to the eukaryotes known from the Proterozoic. Total biomass, and thus organic matter, may have been much smaller. Given the modest amounts of organic matter in Archean sediments and the general estimates of low oxygen pressure in the atmosphere at that time, we can speculate that the balance between production and degradation was such that much of the organic matter produced was buried. The distribution of sedimentary environments of Archean rocks that have been mapped suggests that carbonate shelves were present but perhaps not as widespread. The silica efflux under these conditions would seem to be the same as for the later Proterozoic, although the rates may have differed. Hydrothermal influx was high. If rapid weathering of volcaniclastics dominated the continental surfaces, river silica influx would also have been very high. If the total silica influx were high enough to swamp the sorptive capacity of silicates and organic matter in the ocean, silica concentrations may have risen to levels close to or at saturation with respect to amorphous silica. With high concentrations of silica in seawater, precipitation of amorphous silica in sedimentary layers on the ocean

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floor becomes possibIe. Are the chert layers and iaminae of banded iron formation evidence of this kind of precipnation? The origin of banded iron formations continues to be controversial. Certainly, their properties do not yet allow an unambiguous assessment of the kinds of sedimentary and diagenetic en~ronment in which they were formed. Primary precipitation of iron oxide and silica phases may have been followed by diagenetic segregation. As we learn more about the silica cycle at this early stage of Earth history, the origins of banded iron formations may become clearer. Acknowledgmenis-This work was done with the support of National Science Foundation grants #EAR-8606410 and #EAR-90-17748. I am indebted to many colleagues for discussion of these ideas: A. H. Knoll, R. Buick, H. D. Holland, and over many years, the late Robert M. Garrefs. Knoll, Buick, and A. Basu read the manuscript and made excellent suggestions. Editorial handling: H. C. Helgeson REFERENCES ALLISONC. W., and HILGERTJ. W. ( 1986) Scale microfossils from the Early Cambrian of northwest Canada. J. Paleontol. 60, 9731015. ASMEROMY., JACOBSENS. B., KNOLLA. H., BUTTERFIELDN. J., and SWEETK. ( 1991) Strontium isotope variations of neoproterozoic seawater: Implica~ons for crustal evolution. Geochim. Cusmothim. Acta 55,2883-2894. BARGHOORNE. S. and TYLER S. A. ( 1965 ) Microorganisms from the Gunflint chert. Science 141, 536-577. BASUA. ( I98 1)Weathering before the advent of land plants: Evidence from unaltered detrital K-feldspars in Cambro-Ordovician arenites. Geology 9, 132- 133. BENNETTP. C. ( 1991) Quartz dissolution in organic-rich aqueous systems. Geachim. Cosmochim. Acta 55, 178l-l 797. BENNETTP. C. and SIEGELD. I. ( 1987) Increased solubility of quartz in water due to complexation by dissolved organic compounds. Nature 326,684-687. BERNERR. A. ( 1990) Atmospheric carbon dioxide levels over Phanerozoic time. Science 249: 1382- 1386. BIRNBAUMS. J. and WIREMANJ. W. ( 1984) Bacterial sulfate reduction and pH: Implications for early diagenesis. Chem. Geol. 43, 143-149. BIRNBAUM S. J. and W~REMAN J. W. ( 1985) Sulfate-reducing bacteria and silica ~lubility: A possible mechanism for evaporite diagenesis and silica precipitation in banded iron formations. Canadian J. Ear&hSci. 22, 1904- 1909. CALVERTS. C. ( 1974) Deposition and diagenesis of silica in marine sediments. In Pelagic Sediments: On Land and Under the Sea (ed. K. J. Hsu and H. C. JENKYNS); intl. Assoc.Sedimental. Spec. Pub. 1. p. 273-299. GROTZINGER J. ( 1986) Evolution of early Proterozoic passive-margin carbonate platform, Rocknest Formation, Wopmay orogen, Northwest Territories, Canada. J. Sediment. Petrol. 56,831-847. GROTZINGERJ. and ROYDENL. ( 1990) Elastic strength of the Slave craton at I .9 Gyr and implications for the thermal evolution of the continents. Nature 347,64-t%. HOFFMANP. F. ( 1980) Wopmay orogen: A Wilson cycle of early Proterozoic age in the northwest of the Canadian shield. In The Continental Crust and Its Mineral Deposits (ed. P. W. STRANGWAY); Geol. Assn. Canada Spec. Paper 20, pp. 523-549. HOLLANDH. D. and BEUKESN. J. ( 1990) A paleoweathering profile

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