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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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