Salty Matters

The Blog is written by me, John Warren. Once every three or four weeks or so I will post an article or two on an evaporite topic that has piqued my interest. On the Saltwork Publications webpage (under "the Works") there is a growing library of pdfs and epubs based on these blogs. These articles on the website have much higher resolution extractable graphics in than in the blog. There is also a link to this set of pdfs and epubs on the home page (

Silica mobility and replaced evaporites: 3 - Archean cherts

John Warren - Sunday, August 28, 2016


The two previous articles on silica mobility in evaporitic settings emphasised Phanerozoic examples and discussed silica textures largely tied to the replacement of sulphate evaporite nodules. This article will extend the time frame back to the Archean and also discuss scale controls on massive marine-derived evaporite beds in the early earth. The next article after this focuses on the Proterozoic. In order to extend our discussion into saline Precambrian successions, we must consider changes in ionic proportions and temperatures of the world’s oceans that this involves, and also include the background context of biological evolution of silica-extracting organisms.

Chert deposits clearly preserve a record of secular change in the oceanic silica cycle cross the Precambrian and the Phanerozoic (Maliva et al., 2005), with the chert nodule-evaporite association most obvious in alkaline brine-flushed areas in Phanerozoic sediments (previous 2 articles). Many silicified Phanerozoic evaporite examples co-occur with significant volumes of salts deposited in marine-fed megahalite and megasulphate basins. The evolutionary radiation of silica-secreting organisms across a deep time background is reflected in the transition from abiogenic silica deposition, characteristic of marine and nonmarine settings in the Archean and Proterozoic eons, to the predominantly biologically-controlled marine silica deposits of the Phanerozoic.

Silica levels in the Archean ocean

Estimated silica concentration in Precambrian seawater is 60 ppm SiO2 or more, while silica concentration of much of the modern ocean is controlled by silica-secreting organisms at values of 1 ppm or less to a maximum of 15 ppm (Perry and Lefticariu, 2014). There is no conclusive fossil evidence that such organisms were present in the Precambrian in sufficient abundance to have had a significant influence on the silica cycle, although some later Neoproterozoic protists likely had scales that were siliceous, and Ediacaran sponges certainly produced siliceous spicules. This contrasts with the Phanerozoic, during which the appearance of radiolaria and diatoms changed the locus of silica precipitation (both primary and replacement) from the peritidal and shallow shelf deposits characteristic of the Neoproterozoic, Mesoproterozoic, and much of the Paleoproterozoic, to the deep ocean biogenic deposits since the mid to late Phanerozoic. Comparative petrography of Phanerozoic and Precambrian chert shows an additional early change in nonbiogenic chert deposition occurred toward the end of the Paleoproterozoic era and was marked by the end to widespread primary and early diagenetic silica precipitation in normal marine subtidal environments (Table 1; ca. 1.8 Ga Maliva et al., 2005). Interestingly, the Precambrian transition corresponds to the onset of a plate tectonic regime resembling that of today (Stern, 2007). It was also the time when sulphate levels in the world’s oceans had risen to where gypsum became a primary marine evaporite, as evidenced by large silicified anhydrite nodules (with anhydrite relics) in the late Paleoproterozoic Mallapunyah Fm in the McArthur Basin, Australia (Warren, 2016). Paleoproterozoic early diagenetic “normal marine” cherts generally formed nodules or discontinuous beds within carbonate deposits with similar depositional textures. It seems these “normal marine” cherts formed primarily by carbonate replacement with subsidiary direct silica precipitation. In saline settings cauliflower cherts are also obvious from this time onwards.


Some of these Paleoproterozoic peritidal cherts were associated with iron formations and are distinctly different from younger cherts and appear to have formed largely by direct silica precipitation at or just below the seabed. These primary cherts lack ghosts or inclusions of carbonate precursors, have fine-scale grain fracturing (possibly from syneresis), exhibit low grain-packing densities, and are not associated with unsilicified carbonate deposits of similar depositional composition (Perry and Lefticariu, 2014). Cherts in some Paleoproterozoic iron formations (e.g., the Gunflint Formation, northwestern Lake Superior region) are composed of silica types similar to those in Phanerozoic sinters (e.g., the Devonian Rhynie and Windyfield chert sinters, Scotland, both of which preserved fine-scale cellular detail of Devonian plants, fungi and cyanobacteria, as well as elevated gold levels in the fault feeder system). Such “normal marine cherts lie outside the evaporite focus of this series of articles and for more detail the reader is referred to Perry and Lefticariu, 2014 and references therein.

Archean crustal tectonics and silicification of world-scale evaporites

Archean evaporites were not deposited as saline giants within subsealevel restricted basins created by sialic continent-to-continent proximity setting. In the greenstone terranes that typified the early Archean these tectonic settings simply could not yet exist (Warren, 2016, Chapter 2). Stern (2007) defines plate tectonics as the horizontal motion of Earth’s thermal boundary layer (lithosphere) over the convecting mantle (asthenosphere), and so it is a world-scale system or set of processes mostly driven by lithosphere sinking (subduction pull). He argues that the complete set of processes and metamorphic indicators, associated with modern subduction zones, only became active at the beginning of the Neoproterozoic (≈ 1 Ga). Stern interprets the older record to indicate a progression of tectonic styles from active Archaean tectonics and magmatism (greenstone belts), to something akin to modern plate tectonics at around 1.9 Ga (Figure 1). If so, then modern world-scale plate tectonics only began in the early Neoproterozoic, with the advent of deep subduction zones (blueschists) and associated powerful slab pull mechanisms. Flament et al. (2008) argue that the world’s continents were mostly flooded (mostly covered with shallow ocean waters) until the end of the Archaean and that only 2–3 % of the Earth’s area consisted of emerged continental crust by around 2.5 Ga (aka “water-world”).

It is very likely that the Archaean Earth’s surface was broken up into many smaller plates with volcanic islands and arcs in great abundance (greenstone terranes). Small protocontinents (cratons) formed as crustal rock was melted and remelted by hot spots and recycled in subduction zones. There were no large continents in the Early Archaean, and small protocontinents were probably the norm by the MesoArchaean, when the higher rate of geologic activity (hotter core and mantle) prevented crustal segregations from coalescing into larger units (Figures 1 and 3 ). During the Early-Middle Archaean, Earth’s heat flow was almost three times higher than it is today, because of the greater concentration of radioactive isotopes and the residual heat from the Earth’s accretion, hence the higher ocean temperatures (Figure 2; Eriksson et al. 2004). At that time of a younger cooling earth there was considerably greater tectonic and volcanic activity; the mantle was more fluid and the crust much thinner. This resulted in rapid formation of oceanic crust at ridges and hot spots, and rapid recycling of oceanic crust at subduction zones with oceanic water cycling through hydrothermally active zones somewhat more intensely than today (Zegers and van Keken 2001; Ernst 2009; Flament et al. 2008).

In the Pilbara craton region of Australia significant crustal-scale delamination occurred ≈ 3.49 Ga, just before the production of voluminous TTG (tonalite, trondhjemite, and granodiorite) melts between 3.48 and 3.42 Ga and the accumulation sonic evaporites (Figure 3; Zegers and van Keken 2001). Delamination resulted in rapid uplift, extension, and voluminous magmatism, which are all features of the 3.48–3.42 Ga Pilbara succession. As the delaminated portion was replaced by hot, depleted mantle, melts were produced by both decompressional melting of the mantle, resulting in high-MgO basalts (this is the Salgash Subgroup in the Pilbara craton), and melting of the gabbroic and amphibolitic lower crust, so producing TTG melts. Partial melting of the protocrust to higher levels can be envisaged as a multistep process in which heat was conducted to higher levels and advection of heat occurs by intrusion of partial melts in subsequently higher levels (indicated by purple arrows in Figure 3). TTG melt products that were first intruded were subsequently metamorphosed and possibly partially melted, as can be inferred from the migmatitic gneisses of the Pilbara. This multistep history explains the complex pattern of U-Pb zircon ages of gneisses and granodiorites found within the Pilbara batholiths and the range in geochemical compositions of the Pilbara TTG suite.

Key to the formation of early Archaean evaporites, which indicate a sodium bicarbonate ocean at that time (see next section), is the observation that crustal delamination and the creation of TTG melts led to up to 2 km of crustal uplift (Figure 3). This would have driven some regions of what were submarine sedimentary systems into suprasealevel positions in the Archean waterworld, so creating the potential for hydrographically-isolated subsealevel marine seepage sumps in those portions of the uplifted crust above the zones of delamination. It also explains the centripetal nature of much shallow marine sedimentation of that time. This is cardinal at the broad tectonic scale when comparing the distribution of Archaean and Phanerozoic evaporites (Warren, 2016). Most Archaean evaporite are remnants that are pervasively silicified and underlain by layered igneous complexes, which were dominant across the greenstone seafloor and are associated with bottom-nucleated baryte beds tied to hydrothermal seeps.

Felsic protocontinents (suprasealevel cratons) hosting silicified evaporite remnants probably formed atop Archaean hot spots from a variety of sources: mafic magma melting more felsic rocks, partial melting of mafic rock, and from the metamorphic alteration of felsic sedimentary rocks. Although the first continents formed during the Archaean, rock of this age makes up only 7% of the world’s current cratons; even allowing for erosion and destruction of past formations, evidence suggests that only 5–40 % of the present volume continental crust formed during the Archaean. 

Archean oceans and silicified sodic evaporites 

Chert styles and occurrences in saline settings across deep time clearly show that we cannot carry Phanerozoic silica mobility models in saline lacustrine or CaSO4 evaporite associations directly across time into the deep Precambrian. Rather, comparisons must be made in a context of the evolution of the earth’s atmosphere and associated ocean chemistry, both of which are in part related to the earth's tectonic evolution.

Levels of early Archaean sulphate in the world ocean were probably less than a few percent of the current levels and probably remained so until the evolution of an oxygen-reducing biota into the Proterozoic (Habicht and Canfield 1996; Kah et al. 2004; Warren, 2016). Grotzinger and Kasting (1993) argue that high levels of atmospheric CO2 meant HCO3/Ca ratios were much higher in the Archaean and the Paleoproterozoic oceans than today. All the calcium in seawater was deposited as marine cement-stones and other alkaline earth precipitates well before bicarbonate was depleted and there was no Ca left over to precipitate as gypsum. The early Archaean waterworld ocean was likely a Na–Cl–HCO3 sea, and not the Na–Cl ocean of today (Kempe and Degens 1985; Maisonneuve 1982). This early Archaean hydrosphere had a chemistry similar to that found in modern soda lakes like Lake Magadi and Lake Natron (pathway I brines) and hence the term “soda-lake oceans.” This rather different marine brine chemistry would have precipitated halite and trona/nahcolite, not halite/gypsum. It probably meant that if gypsum/anhydrite did ever precipitate directly from evaporating Archaean seawater it did so only in minor amounts well after the onset of halite precipitation.


The case for nahcolite (NaHCO3) as a primary evaporite (Figure 4a-d), along with halite, in the 3.42 Ga rocks of the Barberton greenstone belt was first documented by Lowe and Fisher-Worrell,1999), both the nahcolite and the halite are silicified. Beds of these silicified sodic evaporite define 5 types of precipitates: (1) large, pseudohexagonal prismatic crystals as much as 20 cm long that increase in diameter upward; (2) small isolated microscopic pseudohexagonal crystals; (3) small, tapering-upward prismatic crystals as much as 5 cm long; (4) small acicular crystallites forming halos around type 1 crystals; and (5) tightly packed, subvertical crystal aggregates within which individual crystals cannot be distinguished. Measurement of interfacial angles between prism and pinacoid faces on types 1 and 2 crystals show four interfacial angles of about 63° and two of about 53°. The morphologies and interfacial angles of these crystals correspond to those of nahcolite, NaHCO3 (Figure 4e). There is no clear evidence for the presence of gypsum in these beds. Sugitani et al. (2003) reported silicified nahcolite (the high CO2 form of sodium carbonate salts; see Warren, 2016, chapter 2) in ≈ 3.2 Ga rocks in the northern part of the Eastern Pilbara block, Western Australia (Figures 4, 5). Coarse, upward-radiating, silicified evaporite crystals in the ca. 3.47–3.46 Ga Strelley Pool Chert (Lowe 1983) show the same habit, geometry, and environmental setting as silicified nahcolite pseudomorphs in the Kromberg Fm. in the Barberton belt, South Africa, and also probably represent silicified NaHCO3 precipitates (Lowe and Tice 2004). Depositional reconstructions in both regions imply a strong hydrothermal association to the silicification of the evaporites in both regions as do bottom-nucleated baryte layers that define seafloor seeps fed by hydrothermal waters moving up faults (Figure 4f; Nijman et al., 1999; van den Boorn et al., 2007).

The pervasive presence of type 1 brines as ocean waters in the early Archean, along with elevated silica levels in most surface ocean waters, compared to the Phanerozoic, implies a significant portion of Archean cherts may also have had a volcanogenic sodium silicate precursor, much like the silicification seen in the modern African rift valley lakes (Eugster and Jones, 1968 and article 1 in this series of articles on silica mobilisation). So in order to decipher possible evaporite-silicification associations we must include aspects of hydrothermal fluid inherent to the Archean, as well as the likely higher surface temperatures that typified highly reducing (anoxic) waters of the early Archean ocean (Figure 3).

Archean evaporite deposition and silicification

Worldwide, the most widespread Archaean depositional environment, especially in early Archaean greenstone terranes, was the mafic plain environment (Condie 2016; Lowe 1994). In this setting, large volumes of basalt and komatiite were erupted to form widespread mostly submarine mafic plains characteristic by ubiquitous pillow structures in the lava interlayers. A second significant sedimentary environment was a deepwater, nonvolcanic setting, where chemical and biochemical cherts, banded iron formation, and carbonate laminites were deposited. The typical lack of evaporite indications in these mostly deepwater sediments indicates an ongoing lack of hydrologic restriction while the sediments were accumulating (waterworld association). The third association, a greywacke-volcanic association becomes more widespread in later Archaean greenstones, which typically sit stratigraphically atop mafic plain units. This association is composed chiefly of greywackes and interbedded calc-alkaline volcanics, hydrothermal precipitates and, in some shallower parts, silicified evaporites. It was perhaps mostly an island arc system and dominantly more open marine as it typically lacks widespread indicators of former marine evaporites. However, more locally it also preserves fluvial and shallow-marine detrital sediments, that were probably deposited locally in Archaean pull-apart basins, and associated with mineralogically mature sediments (quartzarenites, etc.). These more continental associations typified the shallowest to emergent parts of these continental rifts.

Unlike the other two early Archean  greenstone terranes this third terrane type can in places, such as the Pilbara, be tied to sedimentary indicators of a surfacing seafloor, indicated by particular chert and volcaniclastic layers showing mud cracks, wave ripples, tidalites interbedded with hyaloclastics, vuggy cherts, banded iron formations, carbonates and thick now-dissolved and altered type 1 evaporite masses (breccias), perhaps residues of beds formerly dominated by sodium carbonate and halite salts (Figure 5). The Warrawoona Group, preserves many such silicified examples that retain fine detail of primary textures such as mud cracks, oolites, and evaporite crystal casts and pseudomorphs, all indicating shallow-water to emergent deposition atop the mafic plain. In terms of crystal outlines there few if any casts of possible gypsum crystals, more typically, they indicate bladed pseudo-hexagonal, bottom-nucleated nahcolite, trona and in some instances, halite pseudomorphs (Figure 4).

Depositionally, to acquire the needed high salinities, these cherty evaporite units must have risen, at least locally, to shallow near-sealevel depths and at time become emergent, allowing local hydrographically-isolated lacustrine/rift evaporite subaqueous deposition or precipitation of local seepage drawdown salts. Associated primary-textured carbonate and baryte layers interbedded with the cherts are typically minor, bottom-nucleated baryte textures that may likely indicate hydrothermal vent deposits (Figure 4f; Nijman et al., 1999).

Inherent high solubility of any sodium bicarbonate and/or halite salts in what was a hotter burial system, more strongly influenced by hydrothermal circulation than today, meant most of the original sodic evaporite salts were not preserved, unless silicified in early burial. But their presence as silicified pseudomorphs in less-altered greenschist terranes intercalated with volcanics (Figure 4), such as in the Yilgarn, Pilbara and Kaapvaal cratons, clearly shows two things; (1) at times in the early Archaean waterworld there was sufficient hydrographic restriction to allow marine sodian carbonate and sodian chloride evaporites to form and (2) this marine restriction/seepage inflow was probably driven by ongoing volcanism and associated uplift, with evaporites restricted to particular basinwide stratigraphic indicator levels. In the East Pilbara, the early Archaean evaporite stratigraphic level is the Strelley Pool chert, in the Warrawoona group (Figure 5). This is also the level with some of the earliest indications of cellular life-forms (Wacey 2009).

For the original sodic evaporites, it marks the hydrological transition from open marine seafloor to a restricted hydrographically-isolated marine-fed sump basin, surrounded by granite-cored highs with the required uplift likely driven by delamination at the level of the mantle transition (Figures 1 and 3). Given the intimate association of chemical sediments to volcanism in early Archaean greenstone basins, and the sodium bicarbonate ocean chemistry then, compared to the Phanerozoic evaporite hydrochemistries, we can expect a higher proportion of CO2 volatilisation, a higher boron content (tourmalinites) in early Archaean, and a higher level of silicification.

Is the present the key to the past?

The study of silicified evaporites and associated sediments, formed in the early stages of the Earth’s 3.5 Ga sedimentary record, shows that not only has ocean chemistry evolved (see August 24, 2014 blog), the earth’s lithosphere/ plate tectonic character has also evolved (Eriksson et al. 2013). The further back in time, the less reliable is the application of the current plate tectonic paradigm with its strongly lateral movements of crustal blocks and associated plate-scale evaporite basin controls. Phanerozoic evaporites, and the associated silicified sulphate nodules, define a marine-fed seep system where subsealevel continental rifts and continent-continent collision belts favour the formation of mega-evaporite basins (Warren, 2010). Instead, in a substantial portion of the earlier part of the 2 billion year earth history that is the Archaean, shows early-earth evaporite deposition was favored by hydrographic isolation created by strong vertical movement of earth’s crust related to upwelling mantle plumes and crustal delamination with more intense hydrothermal circulation and silicification. There is still no real consensus as to actual time when plate tectonics, as it operates today, actually began, but there is consensus that the present, in terms of plate tectonics, plate-edge collision and evaporite distribution, is not the key to much of the Archaean (Stern 2007; Rollinson 2007).

Uplift and the local accumulation of sodium carbonate Archean evaporites occurred in a depositional setting that was dominated by volcaniclastics,hydrothermal vents and extensional tectonics. Tectonic patterns in these settings have a strongly vertical flavor. In contrast, Phanerozoic salts formed from marine waters with a NaCl dominance with minor bicarbonate compared to calcium, and located mostly in subsealevel sumps formed at interacting sialic plate margins where the dominant tectonic flavor is driven the lateral movement of plates atop a laterally moving asthenosphere and the relative proportion of vilified salts is lower.

Whatever and wherever the onset of Archaean evaporite deposition, all agree that the mechanisms and aerial proportions world-scale plate tectonics were different in early earth history compared to the Phanerozoic. The current argument as to how different is mostly centred on when earth-scale plate tectonic processes became similar to those of today. Given much higher crustal heat flows, it is likely that hydrographically isolated subsealevel depressions, required to form widespread marine evaporites were more localized in the Archaean than today and were more susceptible to hydrothermal alteration, metamorphism and silicification. Appropriate restricted brine sumps would have tended to occur in magmatically-induced uplift zones atop incipient sialic segregations, with crestal subsealevel grabens, which were hydrographically isolated by their surrounds created by supra-sealevel uplift. Once deposited, the higher heat flow in Archaean crust and mantle would also have meant any volumetrically significant evaporites masses were more rapidly recycled, silicified and replaced via diagenetic and metamorphic processes than today.

Some authors have noted that there are no widespread marine evaporites in the Archaean and in the sense of actual preserved salts, this is true. But when one considers that the Archaean crust was much hotter than today and hydrothermal circulation was more active and pervasive, then widespread burial preservation of the primary salts seems highly unlikely. Even in the Neoproterozoic, lesser volumes of the original salt masses remain (Hay et al. 2006). The lack of preserved salts in earlier Precambrian strata is perhaps more a matter of great age, polycyclic metamorphic alteration and the typical proximity to shallow hydrothermal fluids in emergent evaporite forming regions of the Archean waterworld. However we must also ask if the onset of modern styles of plate tectonics also played a role in the relative absence of preserved saline giants in strata older than 1Ga, In the next article we shall look how cooling and the onset of sialic plate tectonics similar to today, altered the types, styles and distributions of silicified and other evaporite salts as the world's oceans moved toward a chemistry more akin to that of today.



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Silica mobility and replaced evaporites: 1 - Alkaline lakes

John Warren - Saturday, July 02, 2016


In this series of blog articles, I plan to look at silica mobility, along with characteristic marine and nonmarine hydrogeochemistries over time, and how these parameters control chert and quartz precipitates and replacements in hypersaline settings. First, I will do this in modern surface and nearsurface settings in the marine and nonmarine realms, ending this first article with a focus on silica distribution and its precipitates in sulphate-depleted saline alkaline lacustrine sediments. Next we shall attempt a synthesis of controls on silica mobility and precipitation in the sulphate-enriched hypersaline marine surface, subsurface and burial realms, as well as defining relevant atmospheric and seawater chemistry changes across deep geologic time. And finally, we shall look at silica mobilisation in subsurface hydrothermal saline settings. The context for this discussion initially comes from the utility of recognising various silicified structures (including chert nodules) as indicative of typical marine (smooth-walled nodules), sulphate evaporite-enriched (cauliflower chert nodules) or sulphate-depleted alkaline lake (Magadi or crocodile skin chert) deposits (Figure 1).


Silica geochemistry

Modern river water typically contains less than 15 ppm dissolved silica, shallow meteoric groundwater typically has 10-50 ppm dissolved silica, while the modern ocean has between 0.5 and 10 ppm dissolved silica (Bridge and Demicco, 2008). Silica concentration is lowest in the ocean’s surface layer (less than 1 mg/l), and relatively constant at around 10 ppm below the thermocline. Deep saline sodium chloride and calcium chloride groundwaters contain comparatively little dissolved silica (30-80 ppm), relative to their total ionic content, although some saline subsurface waters in shale pore waters that also entrain dissolved organic acids can contain up to 330 ppm dissolved silica. Some of the highest values present in significant volumes of surface water are found in saline alkaline lakes with elevated pH levels. These waters can contain more than 1000 ppm silica in solution (Figure 3). One of the highest known natural water values, some 3970 mg/l of dissolved, is from a cold water spring known as Aqua de Rey, near the town of Mt. Shasta, California. The spring has a temperature around 54°C ana pH around 11.6.

At normal environmental pH, the dissolution-precipitation reaction of quartz,

Si02(s)quartz + H20(1) <-> H4SiO40(aq)

produces non-ionized silicic acid (H4SiO40). Because quartz is not very soluble at 25 °C, this reaction puts only ≈6 ppm silicic acid into solution. Therefore, most of the silicic acid in river water and groundwater is considered to come from the incongruent dissolution of silicate minerals, such as feldspars, during weathering. Non-crystalline amorphous silica gels are considerably more soluble, putting up to 120 ppm silicic acid into solution across the normal pH range (Figure 2a). The solubility of quartz is significantly affected by an increase in temperature, and at 300°C, approximately 600 ppm silicic acid is dissolved in groundwater with normal pH (Figure 2b; Verma, 2000; Fleming and Crerar, 1982). Siliceous sinter precipitates where such hot waters rise to the surface in hot springs and then cools, as at Mt Shasta. In the subsurface, cooling hydrothermal waters drive considerable silica mineral replacement and other cements associated with some types of epithermal and halokinetic ore-deposits (later blog). Rising pH also significantly affects the solubility of quartz, and this mechanism helps explain the elevated silica levels in the waters of many alkaline lakes (Figures 2a, 3a). At pH > 9 (at 25 °C) silicic acid dissociates:

H4Si04(aq) <-> H+(aq) + H3Si04(aq)

With the elevated pH of alkaline water, this reaction is driven to the right, and the solubility of both quartz and amorphous silica is greatly enhanced (Figure 1a). Saline alkaline lake waters with elevated pH, as in Lake Magadi in the African Rift Valley and the Alkali Valley playa brines of the south-west USA, consistently contain more than 1000 ppm dissolved silica as H3Si04(aq) (Figure 3a).


Modern siliceous sediments

Modern siliceous sediments accumulate as biogenic marine oozes, biogenic freshwater-lake deposits, chemical precipitates in alkaline lakes, chemical precipitates in soils (silcrete), and chemical precipitates around subaqueous and subaerial hot springs. Significant volumes of dissolved silica occur in the waters of saline alkaline lakes in the African Rift Valley (e.g. Lake Magadi) and a number of Basin and Range playa lakes in Oregon and California (e.g. Lake Abert, Oregon and Alkali Valley playa in California). Inflows in both regions are leaching highly labile volcanics (Figure 3).

Volumetrically, the most significant accumulations of modern siliceous sediments worldwide are constituted by seafloor deposits dominated by opal-A skeletons of planktonic diatoms. Today, diatoms (bacillariophytal algae with siliceous tests) scavenge virtually all of the silica in fresh to somewhat saline surface waters of the continental lakes and most significantly in open ocean waters of the marine realm (Figure 4). Diatoms arose in the Mesozoic, but have become particularly abundant over the past 30 million years, and their remains now dominate silica deposits of the ocean floor and the biogenic siliceous component of the sediment in many less-saline lakes and playa inflows (Knauth, 2003; Katz et al., 2004). Today, the oceans are everywhere undersaturated with amorphous silica, so diatoms build their shells despite thermodynamics of the dissolved content, so that some 90% of dead diatoms’ tests dissolve before they finally settle to the sea floor (Bridge and Demicco, 2008). However, appreciable thicknesses of diatom oozes can accumulate on the sea floor beneath regions with high productivity of diatoms, namely polar areas and areas of oceanic upwelling where there is a high flux of sinking diatoms, such as off the California coast. Other silica-secreting, single-celled eukaryotic plankton include the heterotrophic radiolarians and silicoflagellates. Radiolarian oozes are common beneath equatorial zones of oceanic upwelling.


Pore waters of marine siliceous oozes remain undersaturated with respect to amorphous silica for some depth below the sediment surface, and diagenetic reorganisation of silica is common in deep-sea siliceous sediments. This involves the slow conversion of opal-A to opal-CT, and eventually to microcrystalline quartz, via a complicated series of pathways involving quartz cementing and replacing the oozes. In turn, this leads to the diagenetic precipitation of significant volumes of smooth-walled chert nodules (Figure 1; Hesse, 1989). Worldwide the sampling of oceanic sediment by the Deep Sea Drilling Project (DSDP) has recovered well-developed chert in cemented layers within otherwise still unconsolidated Eocene siliceous oozes. The oldest sedimentary opal-A preserved in DSDP cores is Cretaceous in age, beyond that, cherts are made up of recrystallized quartz.

Dissolved silica and biogenic silica cycling in mesohaline lakes

Diatoms are a significant photosynthetic group driving the accumulation of amorphous silica in the bottom sediments of many freshwater temperate and saline lakes. When salinity and nutrient levels are appropriate, the diatom population in these lakes characteristically undergo explosive population growth. In the lit water column of temperate relatively-freshwater lakes, diatom blooms typically occur in the Spring and Fall of the year. This is when the water column in temperate lakes turns over, allowing nutrient-rich bottom waters to reach the near-surface lit zone at times when the surface water temperature is high enough to support diatom growth. In more saline lakes, diatom blooms occur whenever meromictic lake waters freshen to salinities appropriate for diatom growth. This leads to pulses of biogenic silica accumulating in the laminar distal (profundal) bottom sediments of these saline stratified lakes. As in the oceans, most lake columns in systems with periodic diatom blooms are undersaturated with amorphous silica, so most lacustrine diatom tests dissolve as they sink, and continue to dissolve in the bottom sediment. Closer to the lake shore, the greater volume of silica in the lake bottom sediments tends to come from detrital components washed into the lake as siliciclastic sand and mud.

Chert deposits through time preserve a record of secular change in the oceanic silica cycle. The evolutionary radiation of silica-secreting organisms resulted in a transition from abiological silica deposition, characteristic of the Archean and Proterozoic aeons, to the predominantly biologically-controlled silica deposition of the Phanerozoic (Maliva et al., 2005).

Biogenic lacustrine silica

Diatoms also flourish in the fresher water inflow areas of many salt lakes and playas. They are commonplace primary producers in mesohaline and lower penesaline environments with populations expanding across the lake when salinities are suitable. At less favourable time, healthy diatoms are restricted to refugia areas of fresher water springs and ponds. Their buoyant cells, often augmented by chitin threads or colonial adaptions, enable them to keep within the photic zone better than many other halotolerant algae. Some varieties of benthic diatoms live in lake brines with long-term salinities around 120 ‰, while the upper limit for diatom growth is around 180 ‰ (Clavero et al. 2000; Cook and Coleman 2007; Warren, 2016). The most halotolerant diatom taxa in the saltern ponds of Guerrero Negro are; Amphora subacutiuscula, Nitzschia fusiformis (both Amphora taxa), and Entomoneis sp.; all grow well in salinities ranging from 5 to 150 ‰. Three strains of the diatom Pleurosigma strigosum were unable to grow in salinities of less than 50 ‰ and so are true halophilic alga. A similar assemblage of Amphora sp., along with Cocconeis sp. and Nitzschia sp. dominate the high salinity (150 ‰) saltern ponds near Dry Creek, Adelaide, Australia (Cook and Coleman 2007).

Most mesohaline diatom species flourish at times of freshened surface lake waters or in and about perennial seepage and dissolution ponds (posas) about the edges of some salars, where they can be a major component in some lacustrine stromatolites (Figure 5). Several benthic diatom species are conspicuous in building diatomaceous stromatolites in these freshened (refugia) regions of the saline playa system, for example, Mastogloia sp., Nitzschia sp., Amphora sp., Diploneis sp. They function in a manner analogous to that of cyanobacteria in that they produce extracellular gel, are motile, phototropic, can trap and bind sediment, and create surface irregularities in the biolaminate mat (Winsborough et al., 1994).

Many Quaternary saline lakes have experienced significant fluctuations in water level and salinity across their millennial-scale sedimentary histories. For example, some 10,000 years ago Lake Magadi water depths were hundreds of metres above the present-day water levels and the diatomaceous High Magadi Beds (mostly laminites) were deposited. At that time most of the silica in the mesohaline stratified lake resided in sodium silicates deposited as laminites over most of the profundal lake floor. Diatoms flourished in the fresher waters inflow areas tied to deltaic sediments. Similar diatom rich-zones typify the fresher-water inflow areas of the Jordan River where it flows into the Northern end of the Dead Sea (Garber, 1980).

Silica in alkaline surface waters

It seems diatoms are efficient biogenic vectors for dissolved silica removal, not just in the oceans but also in many mesohaline lake settings. Significant diatom populations occur also in many saline lakes, even some hypersaline alkaline ones. They are important sediment contributors in Lake Manyara (Stoffers and Holdship, 1975), Lake Kivu (Degens et al., 1972), Abert Lake (Phillips and Van Denburgh, 1971), and perhaps also in the Great Salt Lake, Utah (Baxter et al., 2005). Silica levels in the water column of most of the lakesare typically very low (Hahl and Handy, 1969). So, in most mesohaline saline lakes the main biogenic form of silica is as amorphous in diatom skeletons, which tend to periodically accumulate in the bottom sediments with pore waters at the lower end of the hypersaline lacustrine salinity range (Warren, 2016, Chapter 9).

Coorong cherts

However, in some mesohaline to penesaline alkaline lakes, such as those of the Coorong region in South Australia, the ability of a diatom test to survive early burial is likely low. The siliceous frustules tend to dissolve and re-precipitate as amorphous inorganic silica. Abiogenic opal-CT precipitates are commonplace in evaporitic carbonate crusts of a number of Coorong Lakes about the edges of alkaline marginal-marine lagoons and lakes of the Coorong District of South Australia, especially those alkaline lakes containing magnesite or hydromagnesite (Peterson and von der Borch, 1965; Warren, 1988, 1990). Silica precipitation happens within the sediment, or just at the sediment surface, as opal gels (colloids) that can contribute up to 6% by weight of the high-magnesium carbonate sediment in the surface crusts of ephemeral saline lakes such as Milne Lake. Subsequent desiccation of the gel (locally known as yoghurt-textured mud) and consequent cracking creates hardened discs and plates of silica-impregnated mud, about 1 cm thick and 10 cm in diameter. About lake edges, these sediment discs and plates tend to erode into intraclast breccias that coat the uppermost massive unit as crust zones. In some lakes these crust breccia are located immediately edgeward of well-developed hemispherical stromatolites (Figure 6a and b; Warren, 1990).

Under seasonal high-pH conditions, the silica source in a Coorong ephemeral lake and its surrounds dissolves, it then re-precipitates as opal-CT as fresher subsurface groundwater with a lower pH seeps into the lake edges, and mixes with the surface lake brines. Measured pH in Milne Lake surface brines is ≈ 9.5 to 10.2. Silica precipitation tends to occur mostly in the periodically subaerial lake edges during times of incipient lake drying and shrinkage, prior to complete lake desiccation (that is abiogenic silica tends to precipitate in the late spring to early summer of the Southern Hemisphere). The initial silica phase impregnating the lake edge carbonate mud is opal-A. Peterson and von der Borch (1965) argued the likely source of the inorganically precipitated amorphous silica was the dissolution of detrital quartz (sand and silt), which is a common detrital component in the early stages of Holocene lake fill. These older Holocene units are now exposed about some lake edges (Warren, 1998, 1990).

However, diatoms do still thrive in many ephemeral Coorong lakes when surface waters have the appropriate levels of salinity and nutrients. They retreat to refugia about fresh water springs and seeps as the lake dries. Even tough common in plankton populations. intact diatom microfossils (siliceous frustules) have not been recognised in most cores from the same lakes. For example, diatom remains are not present in sapropelic muds in North Stromatolite Lake, a modern hydromagnesite-aragonite lake (Warren, 1990; McKirdy et al., 2010). But, within the organic geochemical constituents of the same cores, there are unusual T-shaped, C20 and C25 highly-branched isoprenoids, which are prominent among the aliphatic hydrocarbons in the extracted organic matter (Figure 7; Hayball et al., 1991). These unusual organic components were later recognised as bacillariophycean algal biomarkers (molecular fossils: McKirdy et al., 1995).

Diatoms are not organic-walled, and the silica of their frustules is highly susceptible to dissolution in present-day alkaline pore waters of this and other Coorong lakes. Hence, soon after burial, their cellular organic matter is destined to become part of the amorphous component of the kerogen (Barker, 1992). It is likely that the direct physical evidence for diatoms (viz. their siliceous frustules) is largely dissolved as waters become alkaline in many mesohaline Coorong ephemeral lakes, so that only biomarkers for a diatom source of the inorganically precipitated silica may remain.

Hence, the ultimate source of the inorganic siliceous carbonate breccia that defines the periodically subaerial edges of many ephemeral hypersaline Coorong lakes is likely from readily dissolved amorphous silica of diatom tests, not the much less soluble quartz, which was postulated as the likely silica source by Peterson and von der Borch (1995). Siliceous breccia zones in the edges of the ephemeral Coorong lakes are intimately tied to characteristic tepee expansion features known as extrusion tepees (Figure 6b). These expansion structures in cemented carbonate crusts are related to the desiccation/cementation of precursor gels washed into fractures beneath mobile sheets of colloid  muds ("yoghurts") that wash about the seasonally shrinking lake edge during the late spring to early summer (Kendall and Warren, 1987). No chert nodules are known to occur in the various Coorong Lakes, only siliceous carbonate mud layers and clasts associated with "extrusion" tepees.

Lake Magadi chert

Hypersaline chert is present as nodules, as well as siliceous breccia layers hosted in Pleistocene sediment that crops out landward of the current lake strandline. Precursors to these modern cherts are thought to initially deposit as late Pleistocene sodium silicates across significant portions of the Lake Magadi basin (Figure 8; Eugster, 1969). Today, these cherty precipitates comprise compact, well-indurated layers and nodules in what are otherwise unconsolidated lake sediments, rich in volcanic debris, sodium silicates and diatoms. The chert shows characteristic reticulate shrinkage cracks on the nodule surface giving it the name crocodile-skin or snake-skin chert or Magadi-style chert (Figure 1). The conversion to chert from its sodium silicate precursor is accompanied by many other enigmatic  textural and structural features such as large desiccation polygons, buckling, reticulation, extrusion, casts of mud-cracks and calcite cements.

Trenching in the regions landward of the current Lake Magadi strandline shows chert-rich zones laterally grade into sediments containing subsurface, still unconsolidated layers of sodium silicate gels dominated by the hydrous sodium silicate magadiite (NaSi7013(0H)3-3H20), with lesser amounts of kenyaite NaSi11O20.5(OH)3.H20 and makatite - NaSi2O3(OH)3.H2O (Figure 9a). Magadiite is a highly siliceous phase, running ≈78% SiO2 by weight (Table 1). When fresh, magadiite is white, soft, putty-like and readily deformable, but it dehydrates rapidly on exposure to air to harden irreversibly into fine-grained cherty aggregate. To date, magadiite has been found only in Quaternary alkaline lacustrine settings. In addition to Lake Magadi, Quaternary magadiite occurs in Lake Bogoria and Lake Kitagata, in Lake Chad in western central Africa, in the flats of Alkali Valley playa in Wyoming, and Trinity County in California (Sebag et al., 2001; Ma et al., 2011).


The solubility-equilibrium trends for silica and amorphous silica are similar, with a marked increase in solubility occurring in more alkaline conditions (pH>9; Figures 1a, 9b; Dietzel and Leftofsky-Papst, 2002). In contrast, SiO2 contents at equilibrium with magadiite show a minimum value at a pH around 8.5 and follow a different dissolution pattern to silica. At low pH the concentration of silica in solution increases, as it also does in alkaline solutions at the other end of the pH spectrum. Thus, the concentration of silica in a solution saturated with respect to magadiite, at constant Na content, is lowest in neutral to slightly alkaline solutions. Below pH 5.9, which is the intersection point of the magadiite and amorphous silica curves, magadiite exhibits a higher solubility than amorphous silica. Thus, at pH < 5.9 magadiite will dissolve, while amorphous silica precipitates (Figure 9; Dietzel and Leftofsky-Papst, 2002).

Conditions associated with the precipitation of magadiite from lake brines in Lake Chad, and probably most other soda lake occurrences, including Lake Magadi, require fluctuations in alkalinity or mixing interfaces between alkaline and less alkaline groundwaters (Figure 9b-d). Sebag et al. (2001) list the following conditions as typical of most modern magadiite occurrences; 1) Elevated alkalinity, typically in the lake dry season (pH >9) allow dissolution of silica, followed by lowering of alkalinity in the wet season driving precipitation of silica (Figure 9b, c), 2) High concentrations of dissolved silica (up to 2700 ppm), 3) Incorporation of sodium ions into the silica lattice that precipitates at the time of silica supersaturation (Figure 9d). Depending on the concentration of Na and Si in the brine at the time of precipitation, various sodium silicate minerals will precipitate (Figure 9).

Two general pathways have been proposed to explain the formation of magadiite in silica-rich sodium carbonate brines: a decrease in pH and evaporative concentration. Magadiite can precipitates when dilute inflow waters flow across a dense, sodium carbonate brine layer rich in dissolved silica interface mixing lowers the pH at the chemocline/halocline. In Lake Chad, and in some American examples in Califonia and Oregon, magadiite may have also precipitated by evaporative concentration or by capillary evaporation of saline, alkaline brines at a shallow subsurface water table. Other inferred mechanisms for sodium silicate precipitation include: 1) subsurface mixing of dilute and saline, alkaline groundwaters, 2) a reduction in pH of an alkaline brine resulting from an influx of biogenic or geothermally sourced CO2, and 3) precipitation from interstitial brines. Different sodium silicate minerals may form according to the concentrations of Na and SiO2 in the brine.

Magadiite (sodium silicate) layers and nodules in Lake Magadi weather into cherts and cherty breccia layers and so define Magadi-style cherts, with a characteristic reticulate, cracked or “crocodile-skin” surface created by shrinkage during the transformation from sodium silicate gel to chert nodule (Figure 1; Schubel and Simonson, 1990). In places, the Magadi chert layers preserve laminae of the original sodium silicate precursor. Conversion of magadiite to bedded and nodular chert is thought to take place close to the sediment surface and be related either to 1) the mobilisation and flushing of sodium by dilute waters in these shallow environments or, 2) to spontaneous conversion to chert in slightly deeper brine-saturated zones Intermediate diagenetic products, including kenyaite, amorphous silica and moganite, may form during the transformation to chert (see inset; Icole and Perinet, 1984; Sheppard and Gude, 1986). Both magadiite and the associated cherts have a distinctive trace element signature, unlike most other cherts (Kerrich et al., 2002).


Eugster et al., (1967) proposed that magadiite of the Lake Magadi High Beds was precipitated in the Late Pleistocene water column by diluting silica-rich, sodium carbonate lake brines with fresher waters at the lake chemocline or mid column interface. Mixing lowered the pH, and although the pH change may have been as little as 0.5, a decrease in pH from 10.3 to 9.8 lowers amorphous silica saturation by more than 500 ppm (Figure 2a). Silica solubility changes very little when pH varies below a maximum of 8. Highly alkaline sodium carbonate waters containing abundant SiO2 readily form in the Magadi rift valley via weathering and rapid subsurface hydrolysis of labile volcanic materials. Biogenically produced CO2 can also reduce the pH of the brine and drive magadiite precipitation (Eugster, 1969). Hay (1968) also suggested that simple evaporative concentration of the brine would lead to magadiite precipitation. Independent of freshwater flushing and pH changes, magadiite decomposes thermally in the lab into quartz and calcite at temperatures of 500-700°C (Lagaly et al., 1975).

In the perimeter sediments of Lake Magadi, Eugster (1969) described what he considered to be an impressive syndepositional result of this transformation of sodium silicates to chert, namely shrinkage megapolygons up to 50 m across in a bedded chert host, with bounding upturned chert ridges up to 2 m high. Historically, the megapolygons, extrusion tepees, convolute folds and other syntransformation features were interpreted as recording the shrinkage-induced flow and collapse of the sediments hosting the magadiite gels, as they lost sodium, dehydrated and shrank.

Subsequent work on the same megapolygonal structures by Behr and Röhricht (2000) concluded the megapolygons are not a response to mineralogical transformation, rather they are part of a suite of prelithification seismite structures in soft, siliceous lake sediment of the precursor Lake Magadi. That is, the chert megapolygons are a soft sediment response to intense deformation and local-scale diapirism, as are the numerous pillow-chert mounds, chert extrusives along dykes and fault ramps, horizontal liquefaction slides with breccias, slumps, petees, flows and shear-structures in the magadiite beds (now all preserved in chert at outcrop).

Collapse, liquefaction and extrusion of the pre-lithified siliceous matrix were caused by seismotectonic rift activity in the lake basin, and it activated fault scarplets and large-scale dyke systems. Seismic activity led to liquefaction and other earthquake-induced intrasediment deformation, especially along fault ramps and on tilted blocks. The textures all indicate the chert megapolygons are a form of seismite and do not mean volume changes in the transition from magadiite to chert. After liquefaction and extrusion, the exposed magadiite material solidified via spontaneous crystallisation to chert in an environment that was characterised by highly variable pH and salinity.

So, since the pioneering work of Eugster and others in the 1960s, three sets of diagenetic processes are now thought to be responsible for driving the conversion of magadiite to Magadi-style (crocodile-skin) chert in Lake Magadi and other soda lakes:

(1) Leaching of sodium by dilute surface runoff during weathering of the High Magadi beds, as evidenced by tracing unweathered beds into outcrop and summarised in the chemical transformation (Eugster, 1969; see inset);

NaSi7O13(OH)3.H2O + H+ —> 7SiO2 + Na+ + 5H2O

(2) Spontaneous release of sodium driving the conversion of magadiite to chert, whatever the nature of through-flushing solutions and environments (see inset). In this process sodium is expelled even in the presence of brines; it does not require the fresh water needed for process 1, and perhaps better explains the occurrence of calcite-filled trona casts in cherts and the presence of chert nodules in unweathered magadiite horizons in Lake Magadi (Hay, 1968, 1970).

(3) To this inorganic perspective on the transformation to form chert, Behr (2002) and Behr and Röhricht (2000) added a biological one. Based on field observations and microbiological studies of the cherts in Lake Magadi region, they argue that inorganic cherts are rare at the type locality of Magadi-style cherts. Rather, as inferred in many modern bacterial dolomites, the cherts at Lake Magadi may have been precipitated as amorphous silica via microbial processes and may not have had a sodium silicate precursor. They go on to note that most of the cherts in the Magadi depression are older than the High Magadi Beds and perhaps developed from flat-topped calcareous bioherms of Pleurocapsa, Gloecocapsa, and other coccoid cyanobacteria, along with thinly bedded filamentous microbial mats, stromatolites, bacterial slimes, diatoms, Dascladiacea colonies and other organic matter accumulations. Silicification occurred from a microbially mediated silicasol, via opal-A to opal-C, with final recrystallisation to a chert of quartzine composition. They conclude that metabolic processes of cyanobacteria controlled the pH of the brine and strongly influenced dissolution-precipitation mechanism that created the chert (Figure 1a). Today the debate as to inorganic versus organic origin of cherts in Lake Magadi continues, and is yet to be resolved.

Surdam et al. (1972) listed the following textures as indicators of Magadi-style cherts that have likely evolved from a sodium silicate gel: 1) Preservation of the soft-sediment deformation features of the putty-like magadiite, such as enterolithic folding, lobate nodular protrusions, casts of mudcracks and trona crystals, and extrusion forms; 2) Contraction features, especially the reticulate cracks and polygonal ridges on the surface of the chert, reflecting the loss of volume in the transition (crocodile-skin chert). If the arguments of Behr and Röhricht (2000) are accepted, then only criteria 2) the characteristic shrinkage-related reticulated surface texture (crocodile-skin) should be used to interpret ancient alkaline lake cherts, along with a lack of any indications of calcium sulphate in penecontemporaneous lake sediments.

Given the type-1 hydrogeochemistry needed for highly alkaline continental brines, evaporite minerals likely to be found in association with Magadi-style cherts are the sodium carbonate salts (trona, gaylussite or pirssonite; searlesite) or their pseudomorphs; gypsum and other forms of calcium sulphate are never present in type 1 (trona -precipitating) brines (Figure 10; Hardie and Eugster, 1970). This contrast with the silica replacement mechanisms that occur when gypsum or anhydrite nodules are silicified.

Across the Phanerozoic rock record, ancient examples of crocodile-skin cherts are not common, compared to documented examples of silicified anhydrite nodules (cauliflower chert). Documented examples include: Cambrian alkaline lacustrine sediments in South Australia (White and Youngs, 1980); Eocene Green River sediments in the USA (Eugster and Surdam, 1973), the Middle Devonian in the Orcadian Basin of Scotland (Parnell, 1988) and fluviolacustrine sediments of the Permian Balzano volcanic complex in Italy (Krainer and Spötl, 1998). For all such ancient occurrences of ancient crocodile-skin chert, it should be remembered that the precipitation mechanism in its type area of Lake Magadi is still contentious. That is, Magadi-type chert, is historically interpreted as being diagenetically derived from magadiite, a hydrous sodium silicate precursor deposited from strongly alkaline lake waters. More recent work in Lake Magadi concludes that the same chert, hosted in the same High Magadi Beds is due to chemical decomposition of pyroclastic deposits by alkaline groundwater, and that chert precipitation is strongly influenced by fluctuating levels of biogenic CO2. The numerous deformation features in the High Magadi Beds in this more recent interpretation are unrelated to the mineralogical transformation of magadiite to chert (Behr, 2002).


Across longer time frames than that preserved in the Pleistocene sediments of Lake Magadi, the chemical proportions of various ionic components in seawater are not constant (Figure 10). The varying proportions are intimately related to the evolution of the world’s atmosphere and rates of seafloor spreading (Warren, 2016; Chapter 2). Sulphate (rather than sulphide) only became a significant component in the world’s ocean around 2Ga. Before that, the world’s oceans and its atmosphere lacked significant oxygen, and entrained much higher proportions of CO2 and methane. In the Archean, the world’s oceans were Na-Cl-Ca-HCO3 waters, not the Na-Cl-Mg-SO4 systems of today and trona, along with halite were primary precipitates in marine hypersaline settings. Under that scenario, it is likely that some marine-associated hypersaline cherts were formed via replacement of sodium silicate precursors. In Phanerozoic strata, an ability to separate cauliflower cherts (after gypsum/anhydrite nodules) from crocodile-skin cherts (associated with silicate gels in trona/natron soda lakes) is considered significant in defining marine-fed versus continental saline hydrologies. Hydrochemistry and textures associated with silicification in CaSO4-rich environments is the topic of the next blog article.


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Seawater chemistry (2 of 2): Precambrian evolution of brine proportions

John Warren - Wednesday, August 26, 2015

We saw in the previous Salty Matters article (part 1 of 2) that ionic proportions of major ions in seawater and oceanic salinity have changed through the Phanerozoic and so influenced the make-up of bittern precipitates once the lower salinity salts (carbonates, gypsum and halite) had precipitated. In the Phanerozoic, seawater was dominantly a Na-K-Mg-Ca-Cl (Ca-rich) brine that changed periodically to a Na-K-Mg-Cl-SO4 (SO4-rich) type, as in the modern ocean. This oscillation across 600 million years forces  number of questions, for example, do similar oscillations in ocean chemistry extend back across the Precambrian? How consistent is the chemistry of the world’s oceans since the early Archean? Does the evaporite evidence in Precambrian sediments support a notion of a primordial reducing atmosphere and/or higher levels of bicarbonate in an early Archean ocean?

Some authors postulate that there have been no significant changes in the major ion proportions in seawater and hence the evaporation mineral series for the past 4 Ga (Morse and Mackenzie, 1998). Others assert that the Archean was dominantly a time of little or no atmospheric oxygen and that ocean waters were reducing anoxic fluids and so sulphate levels were low and sulphide levels high in evaporative marine waters (Krupp et al., 1994). Yet others propose that the bicarbonate to calcium ratio was so high in Archean and Palaeoproterozoic seawater compared to today that all the calcium was used up in widespread abiotic marine aragonite and Mg-calcite precipitates (Sumner and Grotzinger, 2000). In this case trona or nahcolite are likely marine evaporites in the early Archean bitterns (see Figure 1 in part 1). Still others have theorised cyclic changes in oceanic chemistry occurred across much of the Precambrian were similar to those of the Phanerozoic. Such changes were perhaps related to changes in styles and rates of sea floor spreading-hydrothermal circulation in midoceanic ridges (Channer et al., 1997) and the development of tonalitic continents (Knauth, 1998). 

Given that the world's oldest known halites occur in the Bitter Springs Formation in the Amadeus Basin of Australia and that they were deposited some 840 Ma, we can only extend a halite chevron inclusion-based study of ocean chemistry back to that time. These brines were sulphate-depleted, while recrystallised halite from the uppermost Neoproterozoic Salt Range Formation (ca. 545 Ma) in Pakistan, contains solitary inclusions indicating SO4-rich brines (Kovalevych et al., 2006). This supports a similar late Neoproterozoic ocean chemistry to today, as do proportions derived from primary fluid inclusions from the Neoproterozoic Ara Formation of Oman (ca. 545 Ma). It seems that  SO4-rich seawater existed during latest Neoproterozoic time. In contrast while recrystallised halite from the somehat older Bitter Springs Formation contains brine inclusions that are entirely Ca-rich, implying ambient basin brines and the mother seawater were Ca-rich some 830-840 Mas. These combined data, supported by the timing of aragonite and calcite seas, as preserved in various marine carbonates, suggest that during the Neoproterozoic, significant oscillations of the chemical composition of marine brines, and seawater occurred over the last 250 million years of the NeoProterozoic, and that the end-members were similar to those of the Phanerozoic oceans. It seems that Ca-rich seawater dominated for a substantial period of Late Precambrian time (more than 200 Ma) from 850 Ma, until some 650 Ma, this was replaced by SO4-rich seawater, returning to Ca-rich seawater at 530 Ma. 

The detail for much of the remaineder of the Precambrian back to 4 Ga is far less precise than when modelling inclusion chemistries based on actual halites. The oldest documented chevron halite is 850Ma and the oldest bedded anhydrite is 1.2Ga, beyond that, only evaporite pseudomorphs are available to study. So, beyond the 850 Ma record established by halite inclusions in the Bitter Springs Fm., can other Precambrian evaporites especially the calcium sulphates with a record that extends back patchily to the Mesoproterozoic, give indirect clues as to a chemical scenario for the world’s paleo-oceans and brine?


Pseudomorphs, especially of halite hoppers, occur in marine rocks as old as Archean, but are far more common, as are the actual salts, in Proterozoic strata (Figure 1; Warren, 2016). Halite or its pseudomorphs characterise areas of widespread marine chemical sedimentation from the Archean to the present. CaSO4 pseudomorph distribution is more enigmatic. In the 1980s and 1990s, the oldest documented CaSO4 pseudomorphs were thought to cm-sized growth-aligned barytes and cherts in 3.45 Ga metasediments in the Pilbara/North Poleregion of Western Australia. They were interpreted as replacing primary bottom-nucleated gypsum (Figure 2; Barley et al., 1979; Lowe, 1983; Buick and Dunlop, 1990). These barytes and cherts occur in volcaniclastics in association with what are possibly the world’s oldest stromatolites (Hofmann et al., 1999; Allwood et al., 2007). Similar growth-aligned baryte crystals, which initially were also interpreted as likely primary gypsum pseudomorphs, occur in the Nondweni greenstones in South Africa, some 3.4 Ga (Wilson and Versfeld, 1994).


Sequences in both regions are now completely silicified or barytised. At the time they were first documented, the recognition of what were considered shallow-water Early Archean gypsum pseudomorphs at North Pole, Pilbara Craton, caused a re-evaluation of models of a totally reducing Archean atmosphere (Dimroth and Kimberley, 1975; Clemmey and Badham, 1982). The presence of free sulphate in surface brines of the Archean world was thought to imply an at least locally oxygenated hydrosphere. Gypsum precipitating in Archean ocean waters also meant calcium levels in the ocean waters were in excess of bicarbonate, as is in the modern oceans. The presence of free-standing gypsum on the seafloor is incompatible with any model of the Early Archean ocean as a “soda lake.”

However, in both the Pilbara and the South African sequences there are no actual calcium sulphate evaporites preserved, only growth-aligned crystal textures, now preserved as baryte or chert. Textures in baryte ore from Frasnian sediments in Chaudfontaine, Belgium, are near identical to those observed at North Pole, Australia. The Belgian barytes are primary shallow subsea-bottom precipitates with no precursor mineral phase (Figure 2 inset; Dejonghe, 1990). Some workers in the Pilbara feel that the growth-aligned Archean baryte in this region is also a primary seafloor precipitate, formed in the vicinity of hydrothermal vents (Vearncombe et al., 1995; Nijman et al., 1999; Runnegar et al., 2001). As such, it is not secondary after gypsum. A similar hydrothermal discharge model has been developed for aligned barytes in the Barberton Greenstone belt (de Ronde et al., 1994, 1996). 

Based on this more recent analysis, levels of Archean sulphate in the world ocean were probably less than a few percent of the current levels and probably remained so until the evolution of a widespread oxygen-producing biota into the Proterozoic (Figures 3, 4; Habicht and Canfield, 1996; Kah et al., 2004). Barium sulphate is highly insoluble in modern oxygenated seawater. To carry large volumes of barium or sulphur (as sulphide) in seawater solution to the precipitation site required anoxic conditions. If the aligned baryte crystals are primary, their formation still requires sulphate to be locally present on the seafloor, at least in the vicinity of the depositional site. A possible source for local sulphate production in the shallow waters that characterised the North Pole site was shortwave ultraviolet photoxidation of volcanic SO2, indicating an inorganic association (Runnegar et al., 2001). Within barytes in the same 3.47-Ga-old barytes there are microscopic sulphides. These sulphide inclusions show a d34S of 11.6‰, possibly indicating microbial sulphate reduction with H2 as electron donor in what was an anoxic seafloor (Canfield et al. 2004; Shen et al., 2009).

According to Nijman et al. (1999) the occurrence of the North Pole baryte in sedimentary mounds atop growth faults meant sulphate was locally derived via boiling of escaping hydrothermal vent waters enriched in Ba, Si and sulphide. As these hydrothermal waters vented beneath marine water columns perhaps 50 metres deep, they boiled or violently degassed. Consequent mixing with normally stratified seawater, caused instantaneous oxidization of sulphide into sulphate that then, on cooling, combined with the Ba to precipitate as growth-aligned baryte crystals on the seafloor. Conflicting notions (replaced gypsum versus primary baryte) mean that at this stage of our understanding, the bedded baryte evidence cannot be reliably used to support an evaporite paragenesis of gypsum and so infer an Archean ocean with ionic proportions similar to those of today.

Archean and Proterozoic distributions of gypsum have been further complicated by the misidentification of primary aragonite splays and pinolitic siderite marbles as gypsum replacements (Warren 2016; Chapter 15). When these misidentifications are removed from the record it is obvious that calcium sulphate precipitating directly from Archean seawater to form widespread beds did not occur, and that precipitation of aragonite as thick crusts on the sea floor was significantly more abundant than during any subsequent time in earth h istory. In contrast to gypsum, halite pseudomorphs are found throughout the Precambrian (Figure 1;e.g. Boulter and Glover, 1986). 

Grotzinger and Kasting (1993) argue that high levels of atmospheric CO2 meant HCO3/Ca ratios were much higher in the Archean and the Palaeoproterozoic oceans than today. All the calcium in seawater was deposited as marine cementstones and other alkaline earth precipitates well before bicarbonate was depleted and there was no Ca left over to precipitate as gypsum. The early Archean ocean was perhaps a Na–Cl–HCO3 sea, and not the Na–Cl ocean of today (Kempe and Degens, 1985; Maisonneuve, 1982). This early Archean hydrosphere had a chemistry similar to that found in modern soda lakes like Lake Magadi and Lake Natron (pathway I brines) and hence the term “soda-lake oceans” (see Figure 1 in part 1) This rather different marine brine chemistry would have precipitated halite and trona/nahcolite, not halite/gypsum. It probably meant that if gypsum did ever precipitate from Archean seawater it did so only in minor amounts well after the onset of halite precipitation. Excessive sodium in the ocean may help explain the ubiquity of stratiform albitites in much of the Archean. They would have formed throughout the marine realm as early diagenetic replacements of labile volcaniclastics/zeolites in volcanogenic/greenstone terranes).

A case for nahcolite (NaHCO3) as a primary evaporite, along with halite, in the 3.42 Ga rocks of the Barberton greenstone belt was documented by Lowe and Fisher-Worrell (1999). Sugitani et al. (2003) reported silicified nahcolite (the high CO2 form of sodium carbonate salts) in ≈3.2 Ga rocks in the northern part of the Eastern Pilbara block, Western Australia. Coarse, upward-radiating, silicified evaporite crystals in the ca. 3.47–3.46 Ga Strelley Pool Chert (Lowe, 1983) show the same habit, geometry, and environmental setting as nahcolite in the Barberton belt and also probably represent silicified NaHCO3 precipitates (Lowe and Tice, 2004).

Marine nahcolite in the 3.5-3.2 Ga sedimentary record is thought to be evidence of surface temperatures around 70±15°C (Figures 3b, c, 4; Lowe and Tice, 2004). Contemporary early Archean nahcolite (NaHCO3) as a primary evaporitic mineral in a very aggressive weathering regime, in the absence of land vegetation, is best explained by a mixed CH4 and CO2 atmospheric greenhouse. CH4/CO2 ratios were <<1 and pCO2 was at least 100-1000 times the present value, perhaps as high as several bars (Kaufman and Xiao, 2003). The formation of large areas of continental crust at 3.2-3.0 Ga, including the Kaapvaal and Pilbara cratons, resulted in the gradual depletion of atmospheric CO2 through weathering and a lack of marine nahcolite since the early Archean. By 2.9-2.7 Ga, declining pCO2 was associated with climatic cooling and siderite-free soils. 

Transitory CH4/CO2 ratios of ~1 may have resulted in the sporadic formation of organic haze from atmospheric CH4, and are reflected in one or more isotopic excursions involving global deposition of abnormally 13C-depleted organic carbon in sediments of this age. Surface temperatures of <60°C after 2.9 Ga may have allowed an increase in the distribution and productivity of oxygenic photosynthetic microbes (and a decrease in sulphur dependent thermophiles). Eventual lowering of newly formed continental blocks by erosion, reduced loss of atmospheric CO2 due to weathering, and continued long-term tectonic recycling of CO2 resulted in rising pCO2 and decreasing CH4/CO2 ratios in the later Archean and eventual re-establishment of a mainly CO2 greenhouse. Similar events may have been repeated in the latest Archean and earliest Proterozoic, but gradually rising production of O2 effectively kept CH4/CO2 ratios to <<1.


By 2.2-2.0 Ga and perhaps as early as 2.5 Ga, reliable examples of pseudomorphs after primary marine-sourced calcium sulphate first appear in the rock record, but aside from the Karelian beds associated with the Lomagundi Event (LE), widespread stratiform sulphate beds of anhydrite do not appear until 1.2 Ga (Figure 5a). Undeniable CaSO4 nodular and lenticular pseudomorphs are widespread in latest NeoArchean of South Africa and Palaeoproterozoic to Mesoproterozoic sediments of the McArthur Basin, Northern Territory, Australia, and in rocks of Great Slave Lake in northern Canada. For example, in the Malapunyah Formation (1.65 Ga) of the Northern Territory, Australia, the outer portions of numerous decimetre to metre-diameter silicified anhydrite nodules still retain outlines of felted anhydrite laths (pers. obs). The oldest reliable sulphate pseudomorphs after anhydrite and gypsum in Australia come from Palaeoproterozoic cherts in the 2.0-2.2 Ga Bartle Member of the Killara Formation, western Australia (Pirajno and Grey, 2002). These cherts locally retain small amounts of anhydrite (verified by XRD, as well as appearing as highly birefringent flecks in thin sections). Other widespread but younger sulphate pseudomorphs occur in the 1.2 Ga Amundsen Basin in the Canadian Arctic Archipelago. Actual CaSO4 beds outcrop in the 1.2 Ga Society Cliff Formation in Baffin and Bylot Islands of the Canadian Archipelago (Kah et al., 2001, 2004). Sulphate evaporite pseudomophs and nodules in all these Neoproterozoic basins are hosted in sedimentary layers up to tens of metres thick and with lateral extents measured in hundreds of square kilometres. All were laid down in shallow marine, coastal, and alluvial environments under an increasingly oxygenated Meso- to Neoproterozoic atmosphere (Jackson et al., 1987; Walker et al., 1977). After passing from the Archean, by the Mesoproterozoic the hydrosphere contained free sulphate and Ca/HCO3 ratios were lower, leading to a decrease in molar-tooth, herringbone and other carbonate textures indicative of widespread inorganic calcium carbonate saturation in shallow oceanic waters (Figure 6). However, oceanic mother brines for these now-widespread calcium-sulphate evaporites were largely H2S rich with only moderate levels of oxygen in the atmosphere until some 800 Ma (Figure 3a).

The work of Kah et al. (2004) shows that prior to 2.2 Ga, when oxygen began to accumulate in the Earth’s atmosphere, sulphate concentrations in the world’s oceans were low, <1 mM and possibly <200 μM (Figure 5). By 0.8 Ga, oxygen and thus sulphate levels had risen significantly. Sulphate levels were between 1.5 and 4.5 mM, or 5–15% of modern values, for more than a billion years after initial oxygenation of the Earth’s biosphere some 2.2-2.4 Ga and mid -ocean depth waters were anoxic for most of that time (Brocks et al., 2005). Marine sulphate concentrations probably remained low, no more than 35% of modern values, for nearly the entire Proterozoic. A significant rise in biospheric oxygen, and thus oceanic sulphate, may not have occurred until the latest Neoproterozoic (0.54 Ga), just before the Cambrian explosion, when sulphate levels may have reached 20.5 mM, or 75% of present day levels. This is a time when thick sulphate platforms first characterised the salt basins of Oman, prior to that most actual calcium sulphate is in the form of nodules or relatively thin beds.

In a refinement of the sulphate model, Bekker and Holland (2012) note that free sulphate bottom-nucleated sulphate evaporites and not just pseudomorphs were present during the Lomagundi Event (2.22 to 2.06 Ga), and then became relatively scarce once more until some 1.2 Ga. For example, there is a 200 m thick stratigraphic interval of sulphate evaporites of Lomagundi-age, preserved in a shallow-water open-marine siliciclastic and carbonate succession (Lower Jatuli informal group) of Karelia, Russia (Morozov et al., 2010). The Lomagundi Event defines the most extreme and longest lasting isotope excursion of carbon in the world’s marine carbonate record. Bedded gypsum pseudomorphs in the Malmani Group some 2.5 Ga (Gandin and Wright, 2007; Eriksson and Warren, 1983) implies that elevated oceanic sulphate levels that typify the Lomagundi Event may have extended a little further back in time, at least locally (Figure 5).

At the same time as the Lomagundi event, the average ferric iron to total iron (expressed as Fe2O3/Fe|Fe2O3|) ratio of shales increased dramatically. At the end of the Lomagundi Event (LE), the first economic sedimentary phosphorites were deposited, and the carbon isotope values of marine carbonates returned to ≈0.0‰VPDB (Figure 2.50). Thereafter marine sulphate evaporites and phosphorites again became scarce, while the average Fe2O3/Fe|Fe2O3| ratio of shales decreased to values intermediate between those of the Archean and Lomagundi-age shales.

In support of this notion of an “oxygen overshoot,” sulphur isotope work by Reuschel et al. (2012) on the 2.1 Ga dolomitic Tulomozero Fm, which entrains abundant CaSO4 pseudomorphs, concluded that there was a minimum level of 2.5 mM sulphate in the world ocean at that time (Figure 5).

Bekker and Holland (2012) argue the short appearance of sulphate evaporites in Logamundi and the other associated events can be regarded as a ca. 200 Ma “glitch” in the gradual oxidation of the atmosphere–ocean system. It was driven by a positive feedback between the rise in atmospheric O2, the oxidation of pyrite in rocks undergoing weathering, a decrease in the pH of soil and ground water, and an increase in the phosphate flux to the oceans. This sequence led to a major increase in the rate of organic matter burial, a rise in atmospheric oxygen, a large increase in the 13C value for marine carbonates, the deposition of marine evaporites containing gypsum and anhydrite, and the formation of the first commercially important phosphorites. The end of the LE was probably brought about by the weathering of sediments deposited during the LE.

In yet another proposal of hydrosphere-atmosphere evolution, Huston and Logan (2004) argue that the presence of relatively abundant bedded sulphate deposits before 3.2 Ga (as the contentious Archean barytes and chert mentioned earlier) and after 1.8 Ga (as CaSO4 salts), and the peak in banded iron formation abundance between 3.2 and 1.8 Ga, and the aqueous geochemistry of sulphur and iron, when taken together suggest that the redox state and the abundances of sulphur and iron in the hydrosphere varied widely during the Archean and Proterozoic. They propose a layered hydrosphere prior to 3.2 Ga in which sulphate was enriched in an upper oceanic layer, whereas the underlying layer was reduced and sulphur-poor. The sulphate was produced by atmospheric photolytic reactions with volcanic gases in a reducing atmosphere. Mixing of the upper and lower water masses allowed the banded barytes to form prior to 3.2 Ga and created an ocean chemistry where nahcolite was a marine evaporite. Between 3.2 and 2.4 Ga, decreasing volcanogenesis and sulphate reduction removed sulphate from the upper layer, producing broadly uniform, reduced, sulphur-poor and iron-rich oceans.

Whatever the origin of the early Archean baryte and chert, around 2.2 - 2.4 Ga, as a result of increasing atmospheric oxygenation, the flux of sulphate into the hydrosphere by oxidative weathering was greatly enhanced, producing layered oceans, with sulphate-enriched, iron-poor surface waters and reduced, sulphur-poor and iron-rich bottom waters. Gypsum evaporites were increasingly likely as marine precipitates. The rate at which this process proceeded varied between basins depending on the size and local environment of the basin. By 1.8 Ga, the hydrosphere was relatively sulphate-rich and iron-poor throughout. Gypsum was now a widespread marine evaporite. Variations in sulphur and iron abundances suggest that the redox state of the oceans was buffered by iron before 2.4 Ga and by sulphur after 1.6 to 1.8 Ga (Figure 1).

Gypsum in combination with halite was the marine evaporite association from then until now. Seawater was predominantly a Na-Cl±SO4 ocean. Neoproterozoic stratiform sulphates along with widespread halokinetic halite, occur in the Bitter Springs Formation of the Amadeus basin, central Australia (0.8 Ga), its equivalents in the Officer Basin, the Callana beds of the Flinders Ranges and the younger Infracambrian salt basins of the Arabian (Persian) Gulf (≈0.545 Ga; Wells, 1980; Cooper, 1991; Mattes and Conway-Morris, 1990; Edgell, 1991).

The transition to calcium sulphate textures in evaporite pseudomorphs mirrors a marked change in the style of marine carbonates that began around 2.2 to 2.3 Ga when herringbone calcite and precipitated carbonate beds become much less common and the precipitation mode shifted from the seafloor to the water column (Figure 6; Sumner and Grotzinger, 1996, 2000). The boundary also corresponds to the “rusting” of the oceans when oxygen levels became high enough to precipitate widespread banded iron deposits on the seafloor. Microdigitate stromatolites cross this boundary with little effect, suggesting the marked decrease in dissolved iron exerted little influence on them.

The relative scarcity of actual Pre-Phanerozoic salts, not pseudomorphs, especially in the Archean has been used by some to argue that conditions were less favourable for widespread evaporite deposition in the early Precambrian (Cloud, 1972). Others, myself included, feel that the relative scarcity of preserved evaporites in older sequences reflects the greater likelihood of fluid flushing, evaporite dissolution and metasomatism in progressively older rocks. It is likely that oceanic calcium-sulphate evaporites were less common in the Archean, and that sodium carbonates mixed with halite were dominant evaporite salts in the seawater-fed saline giants in appropriate tectonic seepage depressions of the Early Archean. But widespread evaporite deposition from sodium-dominated brines did occur throughout the Archean in large drawdown basins isolated from a surface connection with the ocean. A paucity of preserved bedded evaporite salts in the Precambrian reflects an increased probability of partial or complete evaporite dissolution, remobilization and metasomatism with increasing geological age (see meta-evaporite).

In what is an inclusion study of oldest actual halite, Spear et al., (2014) characterised marine brine chemistry using brine inclusions in the 830 Ma salt of the Browne Formation, Officer Basin, Australia (equiv. to Bitter Springs Fm.). It seems that concentrations of the major ions in these inclusions, except K+ and possibly SO42−, fall within the known range of Phanerozoic seawaters. This ananlysis suggests that mid-Neoproterozoic marine sulphate concentrations were lower (≈90%) than modern values. By the terminal Neoproterozoic, fluid inclusions in halite and evaporite mineralogy from the Khewra Salt of Pakistan and the Ara salt in Oman indicate seawater sulphate levels had risen significantly, to 50%-80% of modern concentrations, which parallels increases in atmospheric and oceanic oxygen.


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