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 (www.saltworkconsultants.com).

Silica mobility and replaced evaporites: 3 - Archean cherts

John Warren - Sunday, August 28, 2016

Introduction

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.

References

 

Buick, R., and J. S. R. Dunlop, 1990, Evaporitic sediments of early Archaean age from the Warrawoona Group, North Pole, Western Australia: Sedimentology, v. 37, p. 247-277.

Buick, R., J. R. Thornett, N. J. McNaughton, J. B. Smith, M. E. Barley, and M. Savage, 1995, Record of emergent continental crust ≈3.5 billion years ago in the Pilbara Region: Nature, v. 375, p. 574 - 777.

Condie, K. C., 2016, Earth as an Evolving Planetary System (3rd edition), Elsevier, 350 p.

Eriksson, P. G., W. Altermann, D. R. Nelson, W. U. Mueller, and O. Catuneanu, 2004, The Precambrian Earth - Tempos and Events: Developments in Precambrian Geology, Elsevier, 941 p.

Eugster, H. P., and B. F. Jones, 1968, Gels Composed of Sodium-Aluminum Silicate, Lake Magadi, Kenya: Science, v. 161, p. 160-163.

Flament, N., N. Coltice, and P. F. Rey, 2008, A case for late-Archaean continental emergence from thermal evolution models and hypsometry: Earth and Planetary Science Letters, v. 275, p. 326-336.

Grotzinger, J. P., and J. F. Kasting, 1993, New constraints on Precambrian ocean composition: Journal of Geology, v. 101, p. 235-243.

Habicht, K. S., and D. E. Canfield, 1996, Sulphur isotope fractionation in modern microbial mats and the evolution of the sulphur cycle: Nature, v. 382, p. 342-343.

Hay, W. W., A. Migdisov, A. N. Balukhovsky, C. N. Wold, S. Flogel, and E. Soding, 2006, Evaporites and the salinity of the ocean during the Phanerozoic: Implications for climate, ocean circulation and life: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 240, p. 3-46.

Kah, L. C., J. K. Bartley, T. D. Frank, and T. W. Lyons, 2006, Reconstructing sea-level change from the internal architecture of stromatolite reefs: an example from the Mesoproterozoic Sulky Formation, Dismal Lakes Group, arctic Canada: Canadian Journal of Earth Sciences, v. 43, p. 653-669.

Kempe, S., and E. T. Degens, 1985, An early soda ocean?: Chemical Geology, v. 53, p. 95-108.

Lowe, D. R., 1983, Restricted shallow-water sedimentation of early Archean stromatolitic and evaporitic strata of the Strelley Pool Chert, Pilbara Block, Western Australia: Precambrian Research, v. 19, p. 239-283.

Lowe, D. R., 1994, Archean greenstone-related sedimentary rocks, in K. C. Condie, ed., Archean Crustal Evolution: Amsterdam, Elsevier, p. 121-170.

Lowe, D. R., and G. Fisher-Worrell, 1999, Sedimentology, mineralogy, and implications of silicified evaporites in the Kromberg Formation, Barberton Greenstone Belt, South Africa, in D. R. Lowe, and G. R. Byerly, eds., Geologic evolution of the Barberton Greenstone Belt, South Africa, Geological Society of America Special Paper, v. 329, p. 167-188.

Lowe, D. R., and M. M. Tice, 2004, Geologic evidence for Archean atmospheric and climatic evolution: Fluctuating levels of CO2, CH4, and O2 with an overriding tectonic control: Geology, v. 32, p. 493-496.

Maisonneuve, J., 1982, The composition of the Precambrian ocean waters: Sedimentary Geology, v. 31, p. 1-11.

Maliva, R. G., A. H. Knoll, and B. M. Simonson, 2005, Secular change in the Precambrian silica cycle: Insights from chert petrology: Geological Society of America Bulletin, v. 117, p. 835-845.

Nijman, W., K. H. de Bruijne, and M. E. Valkering, 1999, Growth fault control of Early Archaean cherts, barite mounds and chert-barite veins, North Pole Dome, Eastern Pilbara, Western Australia: Precambrian Research, v. 95, p. 245-274.

Perry, E. C. J., and L. Lefticariu, 2014, Formation and Geochemistry of Precambrian Cherts, in H. D. Holland, and K. K. Turekian, eds., Treatise on Geochemistry (2nd edition), Elsevier, p. 113-139.

Robert, F., and M. Chaussidon, 2006, A palaeotemperature curve for the Precambrian oceans based on silicon isotopes in cherts: Nature, v. 443 (7114), p. 969-972.

Rollinson, H., 2007, When did plate tectonics begin?: Geology Today, v. 23, p. 186-191.

Stern, R., 2007, When and how did plate tectonics begin? Theoretical and empirical considerations: Chinese Science Bulletin, v. 52, p. 578-591.

Sugitani, K., K. Mimura, K. Suzuki, K. Nagamine, and R. Sugisaki, 2003, Stratigraphy and sedimentary petrology of an Archean volcanic-sedimentary succession at Mt. Goldsworthy in the Pilbara Block, Western Australia: implications of evaporite (nahcolite) and barite deposition: Precambrian Research, v. 120, p. 55-79.

Tänavsuu-Milkeviciene, K., and J. F. Sarg, 2015, Sedimentology of the World Class Organic-Rich Lacustrine System, Piceance Basin, Colorado, in M. E. Smith, and A. R. Carroll, eds., Stratigraphy and Paleolimnology of the Green River Formation, Western USA: New York, Springer, p. 153-182.

van den Boorn, S. H. J. M., M. J. van Bergen, W. Nijman, and P. Z. Vroon, 2007, Dual role of seawater and hydrothermal fluids in Early Archean chert formation: Evidence from silicon isotopes: Geology, v. 35, p. 939-942.

Van Kranendonk, M. J., R. Hugh Smithies, A. H. Hickman, and D. C. Champion, 2007, Review: secular tectonic evolution of Archean continental crust: interplay between horizontal and vertical processes in the formation of the Pilbara Craton, Australia: Terra Nova, v. 19, p. 1-38.

Wacey, D., 2009, Early Life on Earth: A Practical Guide: Topics in Geobiology, 31, Springer.

Warren, J. K., 2010, Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine deposits: Earth-Science Reviews, v. 98, p. 217-268.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Published Feb. 22, 2016: Berlin, Springer, 1854 p.

Zegers, T. E., and P. E. van Keken, 2001, Middle Archean continent formation by crustal delamination: Geology, v. 29, p. 1083-1086.


 

Silica mobility and replaced evaporites: 2 - replaced CaSO4

John Warren - Sunday, July 31, 2016

Postdepositional silicification of sulphate evaporites, that is the precipitation of authigenic silica as a replacement of a CaSO4 host, is the focus of this article, but can be considered a subtopic of the broader styles of silica deposition and silicification that have occurred throughout the geological record from the Precambrian to the Quaternary (Knauth and Epstein 1976; Bustillo 2010). The next article will extend the silica -precipitate discussion back in time across the Proterozoic and into the Archean and consider the influences of atmospheric evolution and seawater chemistry on the styles of silica rocks across deep time. At this point in our discussion, a few relevant geological and mineralogical definitions are needed (see Bustillo, 2010 and Marin-Carbonne et al for more detail). Silica rock is a general term used to define any rock composed mainly of SiO2. In the strict sense, “chert” is used to define a silica rock made primarily of quartz, plus small amounts of opaline minerals, whereas the term “opal” is used to indicate both a mineral and rock. Cherts are sedimentary rocks formed either by direct precipitation from hydrothermal fluids or seawater (known as C-cherts) or by silicification of precursor material (S-cherts). That is, C-cherts are the result of from orthochemical precipitation from seawater (or any Si-rich fluid) and S-cherts are the result of the replacement of a precursor lithology (van den Boorn et al. 2010). This precursor can be evaporitic or volcanogenic sediment (Marin-Carbonne et al., 2014) This article emphasises S-chert examples from the Phanerozoic saline settings, where silica is a secondary phase replacing a pre-exisiting evaporite nodule or crystal. This style of authigenic silica is a common diagenetic constituent in evaporitic carbonates, and occurs in a variety of crystal forms and morphologies (Folk and Pittman 1971; Chowns and Elkins 1974; Knauth 1979; Milliken 1979; Geeslin and Chafetz 1982; Chafetz and Zhang 1998; Scholle and Ulmer-Scholle 2003).

Authigenic silica (S-cherts) can form by: (1) Diagenetic recrystallization of an amorphous silica precursor (Hesse 1989; Knauth 1994); (2) Direct precipitation from aqueous solutions (Mackenzie and Gees 1971; Guidry and Chafetz 2002; Marin et al. 2010); and (3) Direct replacement of pre-existing olcanogenic, carbonate or evaporite host (Hesse 1989; Knauth 1994). Several possible chemical explanations have been suggested to drive the replacement. These include silica precipitation induced by a local decrease in pH that is caused by either biological production of CO2 (Siever 1962), oxidation of sulfide into sulphate (Clayton 1986; Chafetz and Zhang 1998), and mixing of marine and meteoric waters (Knauth 1979).

Types and traits of authigenic silica and cherts

The previous article in this series on silica mobility in evaporitic settings focused on the most mobile (soluble) form of silica known as opal-A or amorphous silica which is defined by a broad peak in XRD determinations (Figure 1). Based on that discussion, it seems there are three main ways modern amorphous silica precipitates; 1) Inorganic precipitate (as in the crusts of the Coorong ephemeral lakes, 2) As a replacement of sodium silicates, such as magadiite (as in alkaline lakes in the African Rift Valley, and 3) Biogenically as in diatom and radiolarian tests in various lakes and the oceans. In all three Opal-A (amorphous opal) is the dominant form of SiO2, but there are other more crystalline forms of sedimentary silica Quaternary sedimentary settings with additional opaline and more quartzose forms. Knauth (1994) classified authigenic silica into, a) 3 types of amorphous opal (opal-A, opal-CT, and opal-C) and, b) 5 types of quartz (granular microcrystalline quartz, megaquartz, length-fast chalcedony, length-slow chalcedony, and zebraic chalcedony). 

Unlike quartz, the opaline minerals are metastable and show different degrees of crystallinity, crystal structure and proportions of water. Jones and Segnit (1971) classified opal minerals into three groups, according to their X-ray diffraction (XRD) patterns (Figure 1a): Opal A (with an XRD pattern that resembles that of amorphous silica), Opal C (which shows four moderately broad peaks that coincide closely with the position of the four most intense peaks of α-cristobalite, plus minor evidence of α-tridymite), and Opal CT (with patterns that show signs of both α-cristobalite and α-tridymite). Opal A can be inorganic, but worldwide is frequently found as siliceous microfossils (diatom frustules, sponge spicules, phytoliths, etc.). Opal C is very rare in sediments. Opal-CT is the most common phase, but its structure can differ owing to its variable water content, the ratio of interlayered cristobalite/tridymite to the amorphous background, and the degree of stacking disorder within the silica framework (Guthrie et al., 1995).


So, amorphous silica is composed of relatively pure SiO2 but with only very local crystallographic order. Amorphous silica includes various kinds of hydrated and dehydrated silica gels, silica glass, siliceous sinter formed in hot springs, and the skeletal materials of silica-secreting organisms. Opal or opaline silica is a solid form of amorphous silica with some included water (Figure 1b). It’s abundant in young cherts, extending back into the Mesozoic. Its geological occurrence is varied it can be by alteration of volcanic ash, precipitation from hot springs, and, volumetrically most significant in the Phanerozoic via precipitation as skeletal material by certain silica-secreting organisms. Opal starts out as what is called opal-A, which shows only a very weak x-ray diffraction pattern, indicating that any crystallographic order is very local. With burial, during the initial stage of diagenesis, opal-A is transformed into opal-CT, which shows a weak x-ray diffraction pattern characteristic of cristobalite). Upon further diagenesis, opal-CT is transformed into crystalline quartz, resulting in chert that consists of an equant mosaic of microquartz crystals. Chalcedony is made up of needles or fibers, often spherulitic, composed of quartz. There’s probably amorphous silica in among the needles, and a variable water content. It is metastable with respect to ordinarily crystalline quartz, but it persists across long time frames; it’s found even in some Paleozoic cherts. Porcellanite is the porous form of chert while silicilyte is a related form that typifies evaporite-associated bacterially-mediated sediment forming a producing reservoir in the South Oman Salt Basin (later blog).

During burial diagenesis, opaline phases age by undergoing successive dissolution-precipitation-recrystallization reactions including the well-known opal A→opal CT→quartz transition (Williams and Crerar, 1985; Williams et al., 1985). These transformations depend mainly on time and temperature, but accelerate in meteoric diagenetic settings, where quartz crystals can form directly, and bypass the opaline silica polymorph phase (Arakel et al., 1989; Bustillo and Alonso-Zarza, 2007). The existence of opal-CT in very young and at-surface rocks (Jones and Renaut, 2007; Jones et al., 1996) shows that time is not necessarily “a cause” in silica diagenesis. According to Bustillo (2010) in continental environments, very rapid silica alteration appears to be related to efficient fluid delivery (i.e., hydrogeology), as much as to time.


When opal-A or opal-CT occur in a sedimentary host, their ageing sets silica free in a dissolved form and so influences the diagenetic evolution of the adjacent carbonates, generally producing silica/carbonate replacements, silica cement, or neoformed silicate clay. Quartz is the last stage of the recrystallization of opals, but can also form directly via replacement or the cementation of voids. Such quartz shows many textures under polarising light. Common quartz can have different crystal sizes and forms crypto-, micro-, meso- or macrocrystalline mosaics. Maliva and Siever (1988) indicated that meso- and macrocrystalline quartz are not produced by ageing but only by direct precipitation during replacement or cementation. Chalcedony is a fibrous-texture quartz made up of several different varieties classified by the orientation of the fibres with respect to the crystal’s c-axis, namely (Figure 2a): Calcedonite (length-fast chalcedony, in which the elongation of the fibres is perpendicular to the crystallographic c-axis), quartzine (length-slow chalcedony, in which the elongation is parallel), lutecite (another type of length-slow chalcedony, in which the fibre axis is inclined by approximately 30°), and helicoidal calcedonite or zebraic chalcedony (which shows a systematic helical twisting of the fibre axes around the crystallographic c-axis). These varieties of chalcedony allow the identification of the environment reigning during the replacement or cementation as acid or non-sulphate (length-fast), or basic or sulphate/magnesium-rich (length-slow) (Figure 2; Folk and Pittman, 1971). The host material, therefore, has geochemical control over the textures of quartz precipitated. Unfortunately, there are exceptions to these rules, and the strict application of these criteria can lead to errors of interpretation.

Moganite is a metastable monoclinic silica polymorph that is structurally similar to quartz (Miehe and Graetsch, 1992). The identification of moganite in the presence of quartz is difficult. It can be detected, however, by detailed XRD analyses with Rietveld refinements, and by other techniques such as Raman and NMR analysis. This mineral is found mixed with quartz in many cherts, preferentially in those that developed in evaporitic environments. However, it can also be produced by the replacement of biogenic carbonates during the interaction of the latter with groundwater (Heaney, 1995). Moganite transforms into quartz, as do the opaline phases, and it probably does so quite readily (Rodgers and Cressey, 2001).

In addition to its replacement style, a number of studies have investigated oxygen isotopic compositions (d18O) in chert to infer climate-driven temperature change through time (Degens and Epstein 1962; Knauth and Epstein 1976; Knauth and Lowe 2003).

 

Silicification of calcium sulphate nodules and isolated crystals

Silicified anhydrite nodules and CaSO4 crystals are widely reported and reliably documented in sediments as old as Paleoproterozoic and as young as Holocene (Table 1). Quartzine and lutecite (aka length-slow chalcedony) typically infill or replace nodules that preserve characteristic cauliflower shapes of the antecedent anhydrite/gypsum nodule (Figure 3; Arbey, 1980; Hesse, 1989). According to Folk and Pittman (1971), rates of nucleation and crystallisation are the primary controls on crystal size and variety of silica precipitating in a void in a dissolving nodule. Rates, in turn, depend on the level of silica saturation or its concentration in the mother brine (Figure 2b, c). According to Keene (1983), precipitation of length-slow quartz is favoured in waters with high SO4 and Mg levels.


High pH levels (alkaline conditions) in the mother solution tend to ionise dissolved silica. Neutral or low pH levels favour silica crystallites made up of combined Si(OH)4 groups. These tend to polymerise into spiral chains at lower pH and higher concentrations. At high concentrations and high pH, the silica precipitates possess a fibrous chalcedonic form reflecting their rapid rates of precipitation. High pH at the precipitation site means silica crystallites also tends to be present in solution as single ionised tetrahedra that attach themselves one by one to the growing surface, so creating fibres of quartz with the c axes oriented parallel to the long axis of the growing fibres (length-slow). Under low pH or in non-sulphate settings the silica is polymerised into spiral silica chains that attach tangentially to the growth surface of the silica gel, with their c-axes parallel to the growing crystal surface and perpendicular to the future direction of the fibres (Figure 2c; length-fast; Folk and Pittman, 1971).

Milliken (1979) summarised the typical petrographic and hand specimen scale features of silica that replaced CaSO4 nodules in Mississippian sediments of southern Kentucky and northern Tennessee (Figure 4). Such nodules typically have knobbly irregular cauliflower-like surfaces, while internal diagnostic textures include: 1) length-slow chalcedony after lathlike evaporites, especially anhydrite; 2) quartzine; and 3) small amounts of lutecite associated either with megaquartz that shows strong undulose extinction, or with euhedral megaquartz (Chowns and Elkins, 1974). The megaquartz often encloses small blebs of residual anhydrite.


Many buried calcium sulphate nodules are silicified in a multistage process that involves both replacement and void filling (West, 1964; Chowns and Elkins, 1974). The process commences about the margins of a nodule (stage 1) with a volume for volume replacement of anhydrite by microcrystalline quartz. It generally ends with the growth of euhedral drusy quartz crystals into a central vug (stage 2 and 3). This mode of replacement exemplifies textural changes as seen from the edge toward the centre of the geode in texture style A in Figure 4. However, as noted by Milliken (1979) this edge inward evolution of the geode or nodule fill is typified by a variety of textural styles, which she denoted a styles A through D.

Stage 1 chalcedony or quartzine mimics or pseudomorphs the felted lath textures of the precursor anhydrite in the outer portion of the nodules in all styles. Anhydrite pseudomorphs occur as radiating or decussate aggregates with a distinctive flow-like pattern indicating a felted anhydrite precursor. Identical decussate and flow textures occur in laths that make up sabkha anhydrite nodules and defines their explosive mode of growth, as well as the typical coalesced nodule texture that, when replaced, ultimately controls the broad-scale “cauliflower” outline of the whole replaced nodule (Figures 3 and 5). And so, as well as silicified lath microtextures seen in thin section, outlines of larger crystals that predated anhydritisation and silicification may be preserved by the nodule margin, these crystal outlines vary from prismatic to bladed. Many silicified nodules still retain the knobbly cauliflower surface morphology of its precursor anhydrite; other nodule edges preserve crystal pseudomorphs with the interfacial outlines of gypsum or anhydrite precursors.


Stage 2 microquartz and quartz fill can assume euhedral faces as they grow into voids created by the dissolution of the nodule. At the same time the quartz may continue to engulf and pseudomorph small areas of residual anhydrite or other less common evaporite salts (e.g. styles A, C, D). Quartz crystals precipitated at this stage are commonly zoned, with more anhydrite inclusions found within the inner region of the pseudomorph. Some quartz crystals are doubly terminated and probably grew via the support of a dissolving meshwork of anhydrite. With the final dissolution of the supporting mesh, these quartz crystals sometimes dropped to the floor of the void to create a geopetal indicator. For example, such highly birefringent anhydrite spots define cauliflower nodules in 2.2 Ga sediments in the Yerrida Basin, Australia (El Tabakh et al., 1999).

Stage 3, the final stage of the void fill is typified by the precipitation of coarse drusy euhedral quartz with no included anhydrite. This coarse quartz resembles coarse vein quartz and often has 18O values indicating temperatures of the mesogenetic or burial realm.

Sometimes the processes of void fill may be arrested to leave a hollow core in the silica-lined geode (Styles A, B, C). The void may be filled later by a different burial stage cement such as baryte, sparry carbonate (e.g. ferroan dolomite or calcite), or even metal sulphides. This is the case with the large (up to 1 m diameter) silicified cauliflower-shaped anhydrite nodules of Proterozoic Malapunyah Formation of the McArthur Basin in Northern Australia where baryte, then metal sulphides and then sparry calcite typify the latter stages of void fill (pers. obs.). Similar fracture-filling baryte characterises the later diagenetic stages of silicified and calcitised anhydrite nodules in the Triassic Bundsandstein redbeds of the Iberian Range of central Spain (Figure 6; Alonso-Zarza et al., 2002). Such geodes are typically excellent indicators of burial cement stratigraphy in a mudstone matrix that otherwise preserves few signs of the evolving pore fluid chemistry. Thus textures and isotopic signatures in a replaced nodule can indicate ongoing diagenesis of the anhydrite nodule that preserves aspects of the shallow active phreatic (eogenetic), the mesogenetic zone with basinal brines and then uplift-related telogenetic fluids.


Internally, cauliflower chert may retain no evidence of former anhydrite lathes mimicked in chalcedony, but can be filled with various styles of coarser-grained megaquartz. The resulting nodules still retain the outline of the precursor evaporite nodule (Figure 3). Work on diagenetic timing of numerous silicified CaSO4 nodules (e.g. Milliken, 1979; Geeslin and Chafetz, 1982; Gao and Land, 1991; Ulmer-Scholle and Scholle, 1994) shows that most silica replacement begins with shallow burial, either in the zone of active phreatic flow or in the upper portion of the zone of compactional flow (probably at depths of less than 500-1000 m). Early silica replacement in the zone of active phreatic flow is indicated by a lack of compressional flattening of the nodule, by the preservation of delicate surface ornamentation and the preservation of compactional drapes around replaced nodules. If replacement of an anhydrite nodule occurs later in the burial cycle, the anhydrite nodule has by then become flattened or sluggy and no longer retain a rugose surface. The result can be a series of “cucumbers” rather than “cauliflowers.”

Milliken’s (1979) isotopic evidence implies much silica replacement in the nodules she studied was relatively early in the burial cycle at temperatures that were < 40°C. Silica was supplied by through flushing pore fluids with compositions ranging from seawater to mixed meteoric-seawater. Of course, nodule replacement by silica or calcite does not have to happen on the way down in the burial cycle; it may also happen during uplift back into the telogenetic realm, where the strata have once again entered the zone of active phreatic flow (Figure 7).

 

Until the turn of the century, there were no documented examples of the process of evaporite replacement by quartz in Quaternary sediments. Now, autochthonous, doubly-terminated, euhedral megaquartz crystals have been observed infilling voids in a gypsum- and anhydrite-bearing Pleistocene sabkha dolomite sequence in the Arabian Gulf, as well as forming overgrowths on detrital quartz grains (Chafetz and Zhang, 1998). These siliceous sabkha precipitates are forming within metres of the present sediment surface with a silica source that is probably recycled biogenic material. Individual quartz crystals attain lengths of 1 mm. Many quartz crystals faces preserve impressions of dolomite rhombs or they partly, or entirely, engulf dolomite rhombohedra. This process of replacement is a response to changing fluid chemistry tied early phreatic burial, to see the full suite of silica replacement textures and the variations in the timing of the replacement means one must study ancient evaporite sequences (Table 1).

  

Overall, the texture of silica infill or replacement in a CaSO4 nodule is dependent on the rate of sulphate dissolution, the timing of silica precipitation and the rate of silica supply. Some nodules are dominated by the early lit-par-lit replacement textures (styles A and C in Figure 4), others have textures indicating silica cement (aligned megaquartz)growing into an open phreatic void left after the complete dissolution of the CaSO4. Such nodules may still retain a hollow centre where the anhydrite once resided (Figure 8). When a silica-filled geode did not start to accumulate silica until after all the CaSO4 dissolved, the primary evidence for an evaporite precursor comes from the shape of the replaced nodule and its stratigraphic position within the evaporitic depositional sequence, e.g. beneath an erosional surface that defines the top of the capillary zone.

  

Not all the anhydrite nodules, now replaced by silica, were syndepositional. Maliva (1987) showed that nodular anhydrite parent, now indicated by quartz geodes in the Sanders Group of Indiana, first precipitated in the subsurface, while its surrounding matrix of normal-marine Sanders Group sediment was still unlithified (Figure 9). Anhydrite nodules formed in the subsurface during early burial as hypersaline reflux brines sank into the normal-marine limestones of the Ramp Creek and Harrodsburg Formations. Silica subsequently replaced the anhydrite nodules. These geodes are almost invariably associated with the development of reflux dolomite.

Similarly, not all silica-replacing anhydrite in a particular region need come from the same source or be emplaced by the same set of processes. Silicified nodules within middle-upper Campanian (Cretaceous) carbonate sediments from the Lafio and Tubilla del Agua sections of the Basque-Cantabrian Basin, northern Spain preserve cauliflower morphologies, together with anhydrite laths enclosed in megaquartz crystals and spherulitic fibrous quartz (quartzine-lutecite). All this shows that they formed by ongoing silica replacement of nodular anhydrite (Figure 10; Gómez-Alday et al., 2002). Anhydrite nodules at Lafio were produced by the percolation of saline marine brines, during a period corresponding to a depositional hiatus. They have d34S and d18O mean values of +18.8‰ and +13.6‰ respectively, consistent with Upper Cretaceous seawater sulphate values. Higher d34S and d18O (mean values of + 21.2‰ and 21.8‰, respectively) characterise nodules in the Tubilla del Agua section and are interpreted as indicating a partial bacterial sulphate reduction process in a more restricted marine environment (Figure 10a). Later calcite replacement and precipitation of geode-filling calcite in the siliceous nodules occurred in both sections, with d13C and d18O values indicating the participation of meteoric waters in both regions (Figure 10b). The synsedimentary activity of the Penacerrada diapir (Kueper salt - Triassic), which lies close to the Lafio section, played a significant role in driving the local shallowing of the basin and in the formation of the silica in the anhydrite nodules. In contrast, eustatic shallowing of the inner marine series in the Tubilla del Agua section led to the generation of morphologically similar quartz geodes, but from waters not influenced by brines derived from the groundwater halo of a diapir.


So far the various papers we have discussed relate the onset of silicification to active phreatic hydrologies (brine reflux or meteoric) typically in evaporites in host rocks that are shallow, either in the early stages of burial or later in the uplift realm. In contrast in a paper discussing silicification of sulphate nodules in Permian (Guadalupian) back-reef carbonates of the Delaware Basin, Ulmer-Scholle et al., 1993, conclude these nodules were silicified in the Mesogenetic realm. Replacement occurred at temperatures of 60-90°C at the same time as hydrocarbons were moving with basinal brines through the adjacent porous matrix (Figure 11). Silicification of these evaporite nodules proceeded from the exterior to the interior of the nodules. The fluid inclusions in the replacive megaquartz are primary, and many contain both hydrocarbons and water. In this setting it seems evaporite silicification was coeval with or slightly postdated hydrocarbon migration and the silica was likely sourced by dissolution of siliciclastics in nearby back-reef units.
 
 

Birnbaum and Wireman (1985) argued that bacterial degradation of organic matter must be important in forming silica precipitates in most evaporites. They demonstrated, through experiment, the strong influence of bacterial sulphate reduction on silica solubility. The ability of sulphate-reducing bacteria to remove silica from solution is related to local changes in pH and hydrogen bonding within amorphous silica, followed by polymerization to higher weight molecules. During silica replacement of sulphate evaporites at relatively shallow burial depths, the pore fluid becomes depleted in dissolved sulphate as it is reduced to H2S by the action of anaerobic sulphate-reducing bacteria, which metabolise sulphate from an anhydrite or gypsum substrate. Where this selective dissolution of the sulphate occurs in the presence of amorphous silica, the reaction is accompanied by the precipitation of silica. Hence the microscale mimicry of the lath outlines in the outer parts of many replaced nodules. According to Birnbaum and Wireman, it reflects bacterially-mediated silica replacement of nodules in relatively shallow burial settings where bacteria flourish.

In summary, in terms of processes and diagenetic settings associated with Phanerozoic evaporite silicification it seems abiological processes, including thermochemical sulphate reduction and hydrocarbon migration, are more important at greater burial depths where bacteria no longer survive. Providing matrix permeability is retained, silica replacement can continue into the thermobaric stage and if the sulphate nodule survives mesogenetic replacement can even persist into exhumation. Replacement under a thermobaric regime is frequently indicated by the preservation of hydrocarbon inclusions in the infilling silica cement. Both BSR and TSR will be discussed further in the next blog article, dealing with silicification associated with ancient evaporites, but with more emphasis on possible hydrochemical contrasts between the Precambrian and Phanerozoic subsurface waters.

References

Alonso-Zarza, A. M., Y. Sánchez-Moya, M. A. Bustillo, A. Sopeña, and A. Delgado, 2002, Silicification and dolomitization of anhydrite nodules in argillaceous terrestrial deposits: an example of meteoric-dominated diagenesis from the Triassic of central Spain: Sedimentology, v. 49, p. 303-317.

Arakel, A. V., G. Jacobson, M. Salehi, and C. M. Hill, 1989, Silicification of calcrete in paleodrainage basins of the Australian arid zone: Australian Journal of Earth Sciences, v. 36, p. 73-89.

Arbey, N., 1980, Les formes de la silice et l’identification des évaporites dans les formations silicifiés: Bulletin, Centre Recherche Exploration–Production Elf- Aquitaine, v. 4, p. 309-365.

Birnbaum, S. J., and J. W. Wireman, 1985, Sulfate-reducing bacteria and silica solubility; a possible mechanism for evaporite diagenesis and silica precipitation in banded iron formations: Canadian Journal of Earth Sciences, v. 22, p. 1904-1909.

Bustillo, M. A., 2010, Chapter 3 Silicification of Continental Carbonates, in A. M. Alonso-Zarza, and L. H. Tanner, eds., Developments in Sedimentology, v. Volume 62, Elsevier, p. 153-178.

Bustillo, M. A., and A. Alonso-Zarza, 2007, Overlapping of pedogenesis and meteoric diagenesis in distal alluvial and shallow lacustrine deposits in the Madrid Basin, Spain: Sedimentary Geology, v. 198, p. 255-271.

Chafetz, H. S., and J. L. Zhang, 1998, Authigenic euhedral megaquartz crystals in a Quaternary dolomite: Journal of Sedimentary Research Section A-Sedimentary Petrology & Processes, v. 68, p. 994-1000.

Chowns, T. M., and J. E. Elkins, 1974, The origin of quartz geodes and cauliflower cherts through the silicification of anhydrite nodules: Journal Sedimentary Petrology, v. 44, p. 885-903.

Clayton, C. J., 1986, The chemical environment of flint formation in Upper Cretaceous chalks, in G. d. G. Sieveking, and M. B. Hart, eds., The Scientific Study of Flint and Chert: Cambridge, Cambridge University Press, p. 43-54.

Degens, E. T., and S. Epstein, 1962, Relationship between O18/O16 ratios in coexisting carbonates, cherts, and diatomites: Bulletin American Association Petroleum Geologists, v. 46, p. 534-542.

El Khoriby, M., 2005, Origin of the gypsum-rich silica nodules, Moghra Formation, Northwest Qattara depression, Western Desert, Egypt: Sedimentary Geology, v. 177, p. 41-55.

El Tabakh, M., K. Grey, F. Pirajno, and B. C. Schreiber, 1999, Pseudomorphs after evaporitic minerals interbedded with 2.2 Ga stromatolites of the Yerrida basin, Western Australia: Origin and significance: Geology, v. 27, p. 871-874.

Eugster, H. P., 1967, Hydrous sodium silicate from Lake Magadi, Kenya: precursors of bedded chert: Science, v. 157, p. 1177-1180.

Folk, R. L., and J. S. Pittman, 1971, Length-slow chalcedony; a new testament for vanished evaporites: Journal Sedimentary Petrology, v. 41, p. 1045-1058.

Gao, G., and L. S. Land, 1991, Nodular chert from the Arbuckle Group, Slick Hills, SW Oklahoma: a combined field, petrographic and isotopic study: Sedimentology, v. 38, p. 857-870.

Geeslin, J. H., and H. S. Chafetz, 1982, Ordovician Aleman ribbon cherts; an example of silicification prior to carbonate lithification: Journal of Sedimentary Petrology, v. 52, p. 1283-1293.

Goldberg, K., S. Morad, I. S. Al-Aasm, and L. F. De Ros, 2011, Diagenesis of Paleozoic playa-lake and ephemeral-stream deposits from the Pimenta Bueno Formation, Siluro-Devonian (?) of the Parecis Basin, central Brazil: Journal of South American Earth Sciences, v. 32, p. 58-74.

Gómez-Alday, J. J., F. Garcia-Garmilla, and J. Elorza, 2002, Origin of quartz geodes from Lano and Tubilla del Agua sections (middle-upper Campanian, Basque-Cantabrian Basin, northern Spain): isotopic differences during diagenetic processes: Geological Journal, v. 37, p. 117-134.

Guidry, S. A., and H. S. Chafetz, 2002, Factors governing subaqueous siliceous sinter precipitation in hot springs: examples from Yellowstone National Park, USA: Sedimentology, v. 49, p. 1253-1267.

Guthrie, G. D., D. Bish, and R. C. Reynolds, 1995, Modeling the X-ray diffraction pattern of opal CT: American Mineralogist, v. 80, p. 869-872.

Hay, R. L., 1968, Chert and its sodium-silicate precursors in sodium-carbonate lakes of east Africa: Contributions to Mineralogy and Petrology, v. 17, p. 255-274.

Hay, R. L., and T. K. Kyser, 2001, Chemical sedimentology and paleoenvironmental history of Lake Olduvai, a Pliocene lake in northern Tanzania: Geological Society of America Bulletin, v. 113, p. 1510-1521.

Heaney, P. J., 1995, Moganite as an indicator for vanished evaporites: a testament reborn?: Journal of Sedimentary Research A: Sedimentary Petrology & Processes, v. A65, p. 633-638.

Henchiri, M., and N. Slim-S'Himi, 2006, Silicification of sulphate evaporites and their carbonate replacements in Eocene marine sediments, Tunisia: two diagenetic trends: Sedimentology, v. 53, p. 1135-1159.

Hesse, R., 1989, Silica diagenesis: origin of inorganic and replacement cherts.: Earth Science Reviews, v. 26, p. 253-284.

Jones, B., and R. W. Renaut, 2007, Microstructural changes accompanying the opal-A to opal-CT transformation: new evidence from the siliceous sinters of Geysir, Haukadalur, Iceland: Sedimentology v. 54, p. 921-949.

Jones, B., R. W. Renaut, and M. R. Rosen, 1996, High-temperature (W901C) calcite precipitation at Waikite Hot Springs, North Island, New Zealand: Journal of the Geological Society of London, v. 153, p. 481-496.

Jones, B. F., S. L. Rettig, and H. P. Eugster, 1967, Silica in alkaline brines: Science, v. 158, p. 1310-1314.

Jones, J. B., and E. R. Segnit, 1971, The nature of opal. Part 1: Nomenclature and constituent phases: Journal of the Geological Society of Australia v. 18, p. 57-68.

Keene, J. B., 1983, Chalcedonic quartz and occurrence of quartzine (length-slow chalcedony) in pelagic sediments: Sedimentology, v. 30, p. 449-454.

Knauth, L. P., 1979, A model for the origin of chert in limestone: Geology, v. 7, p. 274-277.

Knauth, L. P., 1994, Petrogenesis of chert: Reviews in Mineralogy and Geochemistry, v. 29, p. 233-258.

Knauth, L. P., and S. Epstein, 1976, Hydrogen and oxygen isotope ratios in nodular and bedded cherts: Geochimica et Cosmochimica Acta, v. 40, p. 1095-1108.

Knauth, L. P., and D. R. Lowe, 2003, High Archean climatic temperature inferred from oxygen isotope geochemistry of cherts in the 3.5 Ga Swaziland Supergroup, South Africa: Geological Society of America Bulletin, v. 115, p. 566-580.

Krainer, K., and C. Spotl, 1998, Abiogenic silica layers within a fluviolacustrine succession, Balzano volcanic complex, Northern Italy - A Permian analogue for Magadi-type cherts: Sedimentology, v. 45, p. 489-505.

Mackenzie, F. T., and R. Gees, 1971, Quartz: synthesis at earth-surface conditions: Science, v. 173, p. 533-535.

Maliva, R. G., 1987, Quartz geodes; early diagenetic silicified anhydrite nodules related to dolomitization: Journal of Sedimentary Petrology, v. 57, p. 1054-1059.

Marin, J., M. Chaussidon, and F. Robert, 2010, Microscale oxygen isotope variations in 1.9 Ga Gunflint cherts: assessments of diagenesis effects and implications for oceanic paleotemperature reconstructions: Geochimica et Cosmochimica Acta, v. 74, p. 1161030.

 

Marin-Carbonne, J., F. Robert, and M. Chaussidon, 2014, The silicon and oxygen isotope compositions of Precambrian cherts: A record of oceanic paleo-temperatures?: Precambrian Research, v. 247, p. 223-234.

 

Miehe, G., and H. Graetsch, 1992, Crystal structure of moganite: a new structure type for silica: European Journal of Mineralogy, v. 4, p. 693-706.

Milliken, K. L., 1979, The silicified evaporite syndrome; two aspects of silicification history of former evaporite nodules from southern Kentucky and northern Tennessee: Journal Sedimentary Petrology, v. 49, p. 245-256.

Muchez, P., P. Vanderhaeghen, H. El Desouky, J. Schneider, A. Boyce, S. Dewaele, and J. Cailteux, 2008, Anhydrite pseudomorphs and the origin of stratiform Cu–Co ores in the Katangan Copperbelt (Democratic Republic of Congo): Mineralium Deposita, v. 43, p. 575-589.

Nagy, Z. R., I. D. Somerville, J. M. Gregg, S. P. Becker, and K. L. Shelton, 2005, Lower Carboniferous peritidal carbonates and associated evaporites adjacent to the Leinster Massif, southeast Irish Midlands: Geological Journal, v. 40, p. 173-192.

Peterson, M. N. A., and C. C. Von der Borch, 1965, Chert: modern inorganic deposition in a carbonate precipitating localitty: Science, v. 149, p. 1501-1503.

Pirajno, F., and K. Grey, 2002, Chert in the Palaeoproterozoic Bartle Member, Killara Formation, Yerrida Basin, Western Australia: a rift-related playa lake and thermal spring environment?: Precambrian Research, v. 113, p. 169-192.

Rodgers, K. A., and G. Cressey, 2001, The occurrence, detection and significance of moganite (SiO2) among some silica sinters: Mineralogical Magazine, v. 65, p. 157-167.

Scholle, P. A., and D. S. Ulmer-Scholle, 2003, A Color Guide to the Petrography of Carbonate Rocks: Grains, textures, porosity, diagenesis, v. 77: Tulsa, Okla, American Association of Petroleum Geologists Memoir, 459 p.

Siever, R., 1962, Silica solubility, 0°–200 °C., and the diagenesis of siliceous sediments: Journal of Geology, v. 70, p. 127-150.

Tucker, M. E., 1976a, Quartz replaced anhydrite nodules ('Bristol Diamonds') from the Triassic of the Bristol District: Geological Magazine, v. 113, p. 569-574.

Tucker, M. E., 1976b, Replaced evaporites from the late Precambrian of Finnmark, Arctic Norway: Sedimentary Geology, v. 16, p. 193-204.

Ulmer-Scholle, D. S., and P. A. Scholle, 1994, Replacement of evaporites within the Permian Park City Formation, Bighorn Basin, Wyoming, USA: Sedimentology, v. 41, p. 1203-1222.

Ulmer-Scholle, D. S., P. A. Scholle, and P. V. Brady, 1993, Silicification of evaporites in Permian (Guadalupian) back-reef carbonates of the Delaware Basin, west Texas and New Mexico: Journal of Sedimentary Petrology, v. 63, p. 955-965.

 

van den Boorn, S. H. J. M., M. J. van Bergen, W. Nijman, and P. Z. Vroon, 2007, Dual role of seawater and hydrothermal fluids in Early Archean chert formation: Evidence from silicon isotopes: Geology, v. 35, p. 939-942.

 

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Published Feb. 22, 2016: Berlin, Springer, 1854 p.

West, I., 1973, Vanished evaporites; significance of strontium minerals: Journal of Sedimentary Petrology, v. 43, p. 278-279.

West, I. M., 1964, Evaporite diagenesis in the lower Purbeck beds of Dorset [with discussion]: Proc. Yorkshire Geol. Soc., v. 34, p. 315-330.

Wheeler, W. H., and D. A. Textoris, 1978, Triassic limestone and chert of playa origin in North Carolina: Journal of Sedimentary Petrology, v. 48, p. 765-776.

Williams, L. A., and D. A. Crerar, 1985, Silica diagenesis, II. General mechanisms: Journal of Sedimentary Petrology, v. 55, p. 312-321.

Williams, L. A., G. Parks, and D. A. Crerar, 1985, Silica diagenesis, I. Solubility controls: Journal of Sedimentary Petrology v. 55, p. 301-311.

 


 

 

 

 

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.

References

Barker, P., 1992, Differential diatom dissolution in Late Quaternary sediments from Lake Manyara, Tanzania: an experimental approach: Journal of Paleolimnology, v. 7, p. 235-251.

Baxter, B., C. Litchfield, K. Sowers, J. Griffith, P. Dassarma, and S. Dassarma, 2005, Microbial Diversity of Great Salt Lake, in N. Gunde-Cimerman, A. Oren, and A. Plemenitaš, eds., Adaptation to Life at High Salt Concentrations in Archaea, Bacteria, and Eukarya: Dordrecht, Netherlands, Springer, p. 9-25.

Behr, H.-J., and C. Röhricht, 2000, Record of seismotectonic events in siliceous cyanobacterial sediments (Magadi cherts), Lake Magadi, Kenya: International Journal of Earth Sciences, v. 89, p. 268-283.

Behr, H. J., 2002, Magadiite and Magadi chert; a critical analysis of the silica sediments in the Lake Magadi Basin, Kenya, in R. W. Renaut, and G. M. Ashley, eds., Sedimentation in continental rifts, v. 73: Tulsa, Society for Sedimentary Geology (SEPM) Special Publication, p. 257-273.

Bridge, J. S., and R. V. Demicco, 2008, Earth Surface Processes, Landforms and Sediment Deposits, Cambridge University Press, 815 p.

Clavero, E., M. Hernandez-Marine, J. O. Grimalt, and F. Garcia-Pichel, 2000, Salinity tolerance of diatoms from Thalassic hypersaline environments: Journal of Phycology, v. 36, p. 1021-1034.

Cook, F. S., and P. S. J. Coleman, 2007, Benthic diatoms in the salinas of the Dry Creek saltfields, South Australia: Hydrobiologia, v. 576, p. 61-68.

Degens, E. T., R. P. Von Herzer, H.-K. Wong, W. E. Deu­ser, and H. W. Jannasch, 1972, Lake Kivu: Structure, chemistry and biology of an East African Rift lake: Geologische Rundschau, v. 62, p. 245-277.

Dietzel, M., and I. Letofsky-Papst, 2002, Stability of magadiite between 20 and 100°C: Clays and Clay Minerals, v. 50, p. 657-666.

Eugster, H. P., 1967, Hydrous sodium silicate from Lake Magadi, Kenya: precursors of bedded chert: Science, v. 157, p. 1177-1180.

Eugster, H. P., 1969, Inorganic bedded cherts from the Magadi area, Kenya: Contributions Mineralogy and Petrology, v. 22, p. 1-31.

Eugster, H. P., and R. Surdam, 1973, Depositional environment of the Green River Formation of Wyoming: A preliminary report: Bulletin Geological Society of America, v. 84, p. 1115-1120.

Fleming, B. A., and D. A. Crerar, 1982, Silicic acid ionization and calculation of silica solubility at elevated temperature and pH application to geothermal fluid processing and reinjection: Geothermics, v. 11, p. 15-29.

Garber, R. A., 1980, The sedimentology of the Dead Sea: PhD thesis, Rennsaler Polytechnic, 170 p.

Hahl, D. C., and A. H. Handy, 1969, Great Salt Lake, Utah: Chemical and physical variation of the brine, 1963-1966. : Utah Geo!. Mineral. Survey, Water Resour. Bull. 12. 33 pp.

Hardie, L. A., 1984, Evaporites: Marine or non-marine?: American Journal of Science, v. 284, p. 193-240.

Hardie, L. A., and H. P. Eugster, 1970, The evolution of closed-basin brines: Mineralogical Society America Special Paper, v. 3, p. 273-290.

Hay, R. L., 1968, Chert and its sodium-silicate precursors in sodium-carbonate lakes of east Africa: Contributions to Mineralogy and Petrology, v. 17, p. 255-274.

Hay, R. L., 1970, Silicate reactions in three lithofacies of a semi arid basin, Olduvai Gorge, Tanzania: Mineralogical Society of America Special Paper, v. 3, p. 237-255.

Hayball, A. J., D. M. MCKirdy, J. K. Warren, C. C. von der Borch, and D. Padley, 1991, Organic facies of Holocene lacustrine carbonates in North Stromatolite Lake, Coorong Region, South Australia. : 15th International Meet-ing on Organic Geochemistry, Manchester, U.K., Programme and Abstracts, p. 19-20.

Hesse, R., 1989, Silica diagenesis: origin of inorganic and replacement cherts.: Earth Science Reviews, v. 26, p. 253-284.

Icole, M., and G. Perinet, 1984, Les silicates sodiques et les milieux evaporitiques carbonates bicarbonates sodiques: une revue: Revue de Geologie Dynamique et de Geographie Physique, v. 25, p. 167-176.

Katz, M. E., Z. V. Finkel, D. Grzebyk, A. H. Knoll, and P. G. Falkowski, 2004, Evolutionary Trajectories and Biogeochemical Impacts of Marine Eukaryotic Phytoplankton: Annual Review of Ecology, Evolution, and Systematics, v. 35, p. 523-556.

Kendall, C. G. S. C., and J. K. Warren, 1987, A review of the origin and setting of tepees and their associated fabrics: Sedimentology, v. 34, p. 1007-1027.

Kerrich, R., R. W. Renaut, and T. Bonli, 2002, Trace-Element Composition of Cherts from Alkaline Lakes in the East African Rift, Sedimentation in Continental Rifts, v. 73, SEPM Society for Sedimentary Geology, p. 275-294.

Knauth, L. P., and D. R. Lowe, 2003, High Archean climatic temperature inferred from oxygen isotope geochemistry of cherts in the 3.5 Ga Swaziland Supergroup, South Africa: Geological Society of America Bulletin, v. 115, p. 566-580.

Krainer, K., and C. Spotl, 1998, Abiogenic silica layers within a fluviolacustrine succession, Balzano volcanic complex, Northern Italy - A Permian analogue for Magadi-type cherts: Sedimentology, v. 45, p. 489-505.

Krauskopf, K. B., 1967, Introduction to Geochemistry, McGraw-Hill, 721 p.

Lagaly, G., K. Beneke, and A. Weiss, 1975, Magadiite and H-magadiite 1. Sodium magadiite and some of its derivatives: American Mineralogist, v. 60.

Ma, L., T. K. Lowenstein, and J. M. Russell, 2011, A Brine Evolution Model and Mineralogy of Chemical Sediments in a Volcanic Crater, Lake Kitagata, Uganda: Aquatic Geochemistry, v. 17, p. 129-140.

Maliva, R. G., A. H. Knoll, and B. M. Simonson, 2005, Secular change in the Precambrian silica cycle: Insights from chert petrology: Geological Society of America Bulletin, v. 117, p. 835-845.

McKirdy, D. M., A. J. Hayball, J. K. Warren, E. D., and C. C. von der Borch, 2010, Organic facies of Holocene carbonates in North Stromatolite Lake, Coorong region, South Australia: Cadernos Laboratorio Xeolóxico de Laxe, v. 35, p. 127-146.

McKirdy, D. M., C. J. Hepplewhite, B. H. Michaelsen, A. G. Mazzoleni, and Y. Bone, 1995, Origin of sapropels in Holocene lakes of the Coorong region, South Australia, in J. O. Grimalt, and C. Dorronsoro, eds., Organic Geochemistry: Developments and Applications to Energy, Climate, Environment and Human History, AIGOA, Donostia-San Sebastián, p. 183-185.

Parnell, J., 1988, Significance of lacustrine cherts for the environment of source-rock deposition in the Orcadian Basin, Scotland: Geological Society, London, Special Publications, v. 40, p. 205-217.

Peterson, M. N. A., and C. C. Von der Borch, 1965, Chert: modern inorganic deposition in a carbonate precipitating localitty: Science, v. 149, p. 1501-1503.

Phillips, K. N., and A. S. Van Denburgh, 1971, Hydrology and geochemistry of Abert, Summer and Goose Lakes, and other closed-basin lakes in South-central Oregon: U.S. Geo!. Survey Prof. Paper 502B. 86 pp.

Schubel, K. A., and B. M. Simonson, 1990, Petrography and diagenesis of cherts of Lake Magadi, Kenya: Journal of Sedimentary Petrology, v. 60, p. 761-776.

Sebag, D., E. P. Verrecchia, S. J. Lee, and A. Durand, 2001, The natural hydrous sodium silicates from the northern bank of Lake Chad: occurrence, petrology and genesis: Sedimentary Geology, v. 139, p. 15-31.

Sheppard, A. P., and A. J. Gude, 1986, Magadi-type chert - A distinctive diagenetic variety from lacustrine deposits, in F. A. Mumpton, ed., Studies in Diagenesis, v. 1578, US Geological Survey Bulletin, p. 335-343.

Surdam, R. C., H. P. Eugster, and R. H. Mariner, 1972, Magadi-type chert in Jurassic and Eocene to Pleistocene rocks Wyoming: Geol. Soc. Am. Bull., v. 83, p. 1739-1752.

Verma, M. P., 2000, Revised quartz solubility temperature dependence equation along the water–vapor saturation curve: Proceedings of the 2000 World Geothermal Congress. Kyushu, Tohoku, Japan, p. 1927-1932.

Warren, J. K., 1988, Sedimentology of Coorong dolomite in the Salt Creek region, South Australia: Carbonates and Evaporites, v. 3, p. 175-199.

Warren, J. K., 1990, Sedimentology and mineralogy of dolomitic Coorong lakes, South Australia: Journal of Sedimentary Petrology, v. 60, p. 843-858.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Published Feb. 22, 2016: Berlin, Springer, 1854 p.

White, A. H., and B. C. Youngs, 1980, Cambrian alkali playa-lacustrine sequence in the northeastern Officer Basin, South Australia.: Journal of Sedimentary Petrology, v. 50, p. 1279 - 1286.

Winsborough, B. M., J. S. Seeler, S. Golubic, R. L. Folk, and B. Maguire, Jr., 1994, Recent fresh-water lacustrine stromatolites, stromatolitic mats and oncoids from northeastern Mexico, in J. Bertrand-Sarfati, and C. Monty, eds., Phanerozoic Stromatolites II: Amsterdam, Kluwer Academic Publishers, p. 71-100.

 


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