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 (

Salt, Oil, Gas & Metals: What Drives the Link?

John Warren - Friday, March 31, 2017


This article is based on a review written for CSPG Reservoirs and presented orally at the AAPG 2016 meeting in Calgary. It endeavours to give an up-to-date synopsis of how and why ancient salty basins tend to contain elevated levels of oil, gas and metals. It begins with an overview of hydrologies, then places ancient salt bodies in their climatic and tectonic context, and lastly looks at industrial associations and predictors. For brevity, details of many regions, deposits and references are summarised as tables, more complete discussion and referencing can be found in appropriate sections in Warren (2016), or feel free to contact me ( or via with requests for more comprehensive documentation of particular examples.

Evaporite styles reflect intrabasin brine hydrology

If we accept the definition of an evaporite as “A salt rock originally precipitated from a saturated surface or near-surface brine in hydrologies driven by solar evaporation,” then the greater volume of saline mineral salts in the earth’s sedimentary realm are the product of solar heating of brine. There are other sets of mineral salts in the depositional and diagenetic realm with the same mineral composition as evaporite salts, but these salts result from cryogenic, hydrothermal or burial re-equilibration processes (Warren, 2016). For a water molecule to escape into the vapour phase in an evaporitic setting, and so increase the salinity of an enduring brine body or pore water in the capillary zone of a sabkha, the water molecule must; 1) absorb heat energy, 2) be located near the liquid surface, 3) be moving in the proper direction, 4) have sufficient energy to overcome liquid-phase intermolecular forces and, 5) pass through the surface tension inter-face (Figure 1).

Simple physics of molecular escape (rate and intensity) during solar evaporation essentially controls the potency of re-maining brine. Bed textures and mineralogies entomb evidence of concentration levels in the brine and its hydrological setting/stability (Figures 2, 3). This makes evaporites excellent gauges of climate present and past. In combination, relative brine density and specific heat capacity of adjacent brine masses control density and thermal stratification in saline brine bodies located at or near the earth’s surface (Figure 2). Subsequent burial alteration controls later textural evolution by interactions with regional shallow and deep phreatic crossflows (Warren, 2016).

Specific heat is the amount of heat needed to raise one gramme of a substance by 1 °C. For a given amount of heat input, a unit volume of hypersaline water will show a greater increase in temperature than a less salty water (Figure 2c). Given the same degree of insolation, this means density-stratified water bodies tend to be heliothermic, with the lower, denser bottom brine layer being warmer than the somewhat less saline, less dense, upper water layer (Figure 2a). The level in any water column where a marked change in temperature occurs is the thermocline, and in a density-stratified brine mass corresponds to the halocline, also termed a chemocline (Figure 3a). The combination of high temperature, high salinity and lower oxygen levels in the lower brine mass in a heliothermal brine system means only a specialised biota can survive there, often with specialised bacterial populations living in waters just above a saline thermocline. A modern example is the purple sulphur-oxidising community flourishing immediately above the halocline in Lake Mahoney, British Columbia. Fluctuating salinity and nutrient levels endemic to many evaporite depositing regions encourage preservation of elevated levels of organic matter in a variety of hypersaline settings past and present (“feast and famine” associations; Warren, 2011).

As any brine concentrates, its density increases (Figure 2b). The overlying water body must be holomictic for bottom-nucleated salts to accumulate across the subaqueous floor of a brine mass and for dense, saturated brine to sink (reflux) into underlying sediments (Figure 3a). Holomixis means a near homogenous distribution of brine density, temperature, and salinity throughout the brine mass, with internal mixing being ongoing and mostly maintained by wind movement. In contrast, a meromictic brine body is internally stratified, with a lower more-saline, denser, warmer water mass separated across a halocline from an upper, less saline, less-dense, cooler water mass. A longterm halocline hinders chemical or physical changes in the underlying denser waters, so shutting down bottom nucleation, as well as slowing and ultimately stopping brine reflux. A permanently stratified system is ectogenic, while a brine column that is temporarily stratified is endogenic.

Over decades, saline water masses can change from ectogenic to endogenic. In February 1979, salinity equalisation drove the mixing of upper and lower water masses in the Dead Sea, resulting in a holomictic water body. Since then, aside from short episodes of storm-flood-driven freshening of the upper water mass in 1980 and 1994 (meromixis), the Dead Sea has been holomictic, and halite has been accumulating on the deep lake floor (Gertman and Hecht, 2002). Before 1979, the Dead Sea had been a stratified system for at least 400 years and only pelagic carbonate laminites with minor gypsum, not halite, accumulated on the deep lake floor, beneath a 370-380 m deep brine column.

Holomixis permits deposition of a coherent salt layer across the whole basin floor, beneath both shallow and deep brine columns. Density stratification allows evaporitic salts to crystallise only in the upper water mass or at the upper brine - lower brine interface, so bottom nucleation tends to occur on the shallower lake floor, where it lies above the halocline (Figure 3a). That is, long-term (ectogenic) column stratification mean bottom nucleation of salts can only occur where the upper salt-saturated brine mass intersects the sediment bottom, with a pelagic settling of salts occurring deeper out in the depositional basin, as in the Dead Sea prior to February 1979 (Figure 3a). The bottom growth of crystals cannot occur on a deep bottom located beneath a density-stratified system, as there is no mechanism to drive ongoing supersaturation in the lower water mass. For the same reason, constant brine reflux driving sinking of a dense brine into sediments beneath the floor of the evaporite basin can only proceed if significant regions of the overlying brine mass are holomictic. Deposition of capillary salts (sabkha deposits) occurs in subaerial settings, wherever the saline capillary zone inter-sects the land surface (Figure 3b).

When salts are accumulating beneath a holomictic brine mass, textures in bottom nucleates is controlled by the stability of the overlying brine column (Figure 3b). When the overlying column is deep (>30-100m) then, other than areas on the deep bottom of local phreatic spring-fed outflows, there is no general hydrochemical mechanism to drive fluctuations in bottom-brine chemistry. The resulting deep bottom precipitates tend to be monomineralogic crystal clusters, possibly encased by re-transported material washed in from the shallower surrounds (Figure 3a, 3b). In contrast, when the overlying brine column is shallow (<30m and typically <5-10m) then the chemistry and stability of the brine varies on a shorter term (daily-weekly) basis, so more layered bi-mineralogic bottom-nucleates can accumulate as layered to laminated salt beds. In addition, all evaporite sediments can be reworked by bottom currents, with similar textures to those that characterise siliciclastic and mechanically-modified carbonate sediments (Figure 3b).

"Now" versus "then" in evaporite deposition

Uniformitarianism is an essential tenet of geological understanding. Yet, when we look at evaporite volumes and depositional settings across deep time, we see that the diversity of modern evaporite analogues is constrained by a deficit in two conditions, specifically; 1) the current lack of greenhouse eustasy; contemporaneous atmospheric conditions and sea levels are controlled by the earth’s current icehouse climate mode and have been for the last 10-12 million years, and 2) the current lack at the plate-edge scale of marine seepage into large hydrographically-isolated oceanic sump basins (Warren, 2010). Both situations circumscribe different hydrologies and eustasies compared to continental-fed in-flows that typify the world's current larger evaporite basins. Today, and across the Quaternary, the largest and thickest salt stacks, with areal extents up to 10,000 km2 and thickness up to 900 m, tend to precipitate in the lower parts of suprasealevel intermontane lacustrine sumps located in tectonically active parts of continental interiors, such as Salar Atacama and Salar di Uyuni in the Andean Altiplano (Figure 4a). Most ancient evaporites are marine-fed and were deposited in huge hydrographically-isolated subsealevel marine-seepage sumps located in intracratonic basins or within rifts or compressional sutures. Often, the areal extents of these ancient systems were more than 250,000 km2, this is more than two orders of magnitude larger than any Quaternary evaporite deposit. Bedded (pre-halokinetic) thicknesses could be more than a kilometre.

Ancient marine saline giants (megahalites and megasulphates) accrued in either of two plate-scale settings, which at times merged into one another, namely; 1) Platform evaporites (Figure 5) and 2) Basinwide evaporites (Figure 6). The first major contrast with nonmarine continental dominance in Quaternary evaporite settings is the fact that platform evaporites require greenhouse eustasy, the second is that basinwide evaporites require tectonically- and hydrographically-isolated widespread subsealevel depressions, typically found along plate edges with continent-continent proximity (Figure 5).

Neither condition is present on the current earth surface. For basinwides, suitable hydrologic conditions were last present during the Messinian Salinity Crisis in the Mediterranean region, and platform evaporite settings were last present on earth across large parts of the Middle East carbonate platform during the Eocene (Tables 1, 2). There is a third group of ancient evaporite deposits; it encompasses all nonmarine lacustrine beds past and present (Table 3). This group has same-scale modern-ancient counterparts, unlike ancient marine platform and basinwide evaporites (Figure 4a; Warren, 2010, 2016).

Platform evaporites

Are made up of stratiform beds, usually <50 m thick and composed of stacked <1 to 5 m thick parasequences or evaporite cycles, with a variably-present restricted-marine carbonate unit at a cycle base (Table 1). Salts were deposited as mixed evaporitic mudflat and saltern evaporites, sometimes with local accumulations of bittern salts. Typically, platform salts were deposited in laterally extensive (>50-100 km wide), hydrographically-isolated, subsealevel marine-seepage lagoons (salterns) or evaporitic mudflats (sabkhas and salinas). These regions had no same-scale modern coun-terparts and extended as widespread depositional sheets across large portions of hydrographically isolated marine platform areas, which passed seaward across a subaerial seepage barrier into open marine sediments (Figure 5). In marine margin epeiric settings, such as the Jurassic Arab/Hith and Permian Khuff cycles of the Middle East or the Cretaceous Ferry Lake Anhydrite in the Gulf of Mexico, these platform evaporites are intercalated with shoalwater marine-influenced carbonate shelf/ramp sediments, which in turn pass basinward across a subaerial sill into open marine carbonates. Landward they pass into arid zone continental siliciclastics or carbonate mudflats.

Platform evaporite deposition occurred in both pericontinental and epicontinental settings, at times of low-amplitude 4th and 5th order sealevel changes, which typify greenhouse eustasy (Figure 5; Warren, 2010). Platform evaporites also typify the saline stages of some intracratonic basins. Platform evaporites cannot form in the high-amplitude, high-frequency sealevel changes of icehouse eustasy. The 100m+ amplitude oscillations of icehouse times mean sealevel falls off the shelf edge every 100,000 years, so any evaporite that had formed on the platform is subaerially exposed and leached. Fourth order high-amplitude icehouse eustatic cycles also tend to prevent laterally-continuous carbonate sediment barriers forming at the top of the shelf to slope break and so icehouse evaporite systems tend not to be hydrographically isolated (drawdown) at the platform scale. Rather icehouse eustasy favours nonmarine evaporites as the dominant style, along with small ephemeral marine-margin salt bodies, as seen today in the bedded Holocene halites and gypsums of Lake Macleod in coastal West Australia.

Ancient platform evaporite successions may contain halite beds, especially in intracratonic basinwide settings, but periplatform settings, outside of intracratonic basins, are typically dominated by 5–40 m thick Ca-sulphate beds intercalated with normal-marine platform carbonates (Table 1). The lateral extent of these epeiric platform sulphate bodies, like the Middle Anhydrite Member of the Permian Khuff Fm. of Saudi Arabia and the UAE, with a current area of more than 1,206,700 km2, constitute some of the most aerially-extensive evaporite beds ever deposited.

Basinwide evaporites

Are made up of thick evaporite units >50–100 m thick made up of varying combinations of deepwater and shallow wa-]ter evaporites (Table 2). They retain textural evidence of different but synchronous local depositional settings, including mudflat, saltern, slope and basin (Figures 6). When basinwide evaporite deposition occurs, the whole basin hydrology is evaporitic, holomictic, and typically saturated with the same mineral phase across vast areas of the basin floor, as in the Dead Sea basin today. The Dead Sea has a more limited lateral scale than ancient basinwides but currently has halite forming simultaneously as; 1) decimeter-thick chevron-dominated beds on the saline-pan floor of the shallow parts around the basin edge in waters typically less than 1-10 metres deep, and 2) as coarse inclusion-poor crystal meshworks of halite on the deep basin floor that sits below a halite-saturated brine column up to hundreds of metres deep (Figure 3a). Ancient basinwide successions are usually dominated by thick massive salt beds, generally more than 100-500 m thick. Deposits are made up of stacked thick halite beds, but can also contain substantial volumes of thick-bedded Ca-sulphate and evaporitic carbonate, as in the intracratonic basinwide accumulations of the Delaware and Otto Fiord Basins (Table 2).

Owing to inherent purity and thickness of the deposited halite, many halite-dominant basinwide beds are also remobilized, via loading or tectonics, into various halokinetic geometries (Hudec and Jackson, 2007). Some basinwide systems (mostly marine-fed intracratonic settings) entrain significant accumulations of marine-fed potash salts, as in the Devoni-an Prairie Evaporite of western Canada. In contrast, all Quaternary examples of commercial potash deposits are accumulating in continental lacustrine systems (Warren 2016; Chapter 11).

Basinwide evaporite deposits are the result of a combination of tectonic and hydrological circumstances that are not currently active on the world’s surface (Figure 4b). They were last active in the Late Miocene (Messinian), in association with soft-suture collision basins tied to the Alpine-Himalaya orogenic belt, and in Middle Miocene (Badenian) basins developed in the early rift stages of the Red Sea. Basinwide systems will be active again in the future at sites and times of appropriate plate-plate interaction, when two continental plate edges are nearby, and the intervening seafloor is in or near a plate-edge rift or suture and is both subsealevel and hydrographically isolated (Figure 6).

Lacustrine (nonmarine) evaporites

Quaternary continental playa/lacustrine are constructed of stratiform salt units, with the greater volume of saline sediment accumulating in lower, more-saline portions of the lacustrine landscape. Beds are usually dominated by nodular gypsum and displacive halite, deposited in extensive evaporitic mudflats and saltpans with textures heavily overprinted by capillary wicking, rather than as bedded bottom-nucleated layers on the subaqueous floors of perennial brine lakes (Figure 3b; Ruch et al., 2012). In ancient counterparts, the total saline lacustrine thickness ranged from meters to hundreds of meters, with lateral extents measured in tens to hundreds of kilometres (Table 3; Figure 4a). Lacustrine salt beds are separated vertically, and usually surrounded by, deposits of lacustrine muds, alluvial fans, ephemeral streams, sheet floods, eolian sands, and redbeds. As today, ancient lacustrine salts accumulated in endorheic or highly restricted discharge basins, with perennial saline water masses tending to occur in the drainage sumps of steep-sided drainage basins (Warren, 2010, 2016). Saline lake basins accumulating gypsum, or more saline salts like halite or glauberite, typically have a shallow water table in peripheral saline mudflat areas and so are dominated by continental sabkha textures. Nearby is the lowermost part of the lacustrine depression or sump where deposition is typified by ephemeral ponded brine pan deposits, rather than permanent saline waters.

Saline lacustrine mineralogies depend on compositions of inflowing waters, so depositional sumps in regions with non-marine ionic proportions in the feeder inflow, accumulate thick sequences of nonmarine bedded salts dominated by trona, glauberite, and thenardite. In contrast, nonmarine areas with thalassic (seawater-like) inflows tend to accumulate more typical sequences of halite, gypsum, and anhydrite.

Across the Quaternary, less-saline perennial saline-lake beds tend to occur during more humid climate periods in the same continental-lacustrine depressions where saline-pan beds form (e.g., Lake Magadi, Great Salt Lake, Lake Urmia). On a smaller scale, in some modern saline lake basins, parts of the lake floor can be permanently located below the water surface (Northern Basin in the Dead Sea or Lake Asal). In some modern saline sumps dominated by mudflats, a perennial saline lake water mass is located toward the edge of a more central salt-flat zone, forming a perennial water filled “moat” facies surrounding a seasonally desiccated saline pan (as in Salar, de Atacama, Salar de Uyuni, Lake Magadi, Lake Natron). These permanent to near-permanent saline water “moat” regions are typically created where fresher inflows encounter saltier beds of the lake centre, dissolve them, and so form water-filled peripheral depressions. Bottom sediment in the moats tend to be mesohaline carbonate laminites, which can contain TOC levels as high as 12%.

High-water stage perennial saline lacustrine sediments tend to be carbonate-rich or silica-rich (diatomaceous) laminites. Ancient examples of large saline lacustrine deposits made up of alternating humid and desiccated lacustrine units include the Eocene Green River Formation of Wyoming and the Permian Pingdiquan Formation of the Junggar Basin, Chi-na (Table 4). Evaporites deposited in a suprasealevel lacustrine basin (especially Neogene deposits) have numerous same-scale Quaternary analogues, unlike the more voluminous ancient marine platform and basinwide evaporites (Figure 4a).

Salt punches above its weight, but why? (facilitator for economic accumulations of oil, gas and metals)

In terms of total mass in sedimentary basins, the proportion of evaporite across the world’s Phanerozoic basins is rarely more than 2% (Figure 7; Ronov, 1980). Today we have comprehensive documentation that salt horizons, their brines, associated dissolution and alteration conduits control significant economic associations of oil, gas and metals (Warren, 2016):

  • 50% of world’s carbonate reservoirs (seals, traps and source rocks)
  • All the world’s supergiant oil and gas fields in thrusts (seals and structural traps).
  • All supergiant sedimentary copper deposits (halokinetic brine focus)
  • 50% of world’s giant SedEx deposits (halokinetic brine focus).
  • 80% of giant MVT deposits (sulphate-fixer & brine interface)
  • World’s largest Phanerozoic Ni deposit (meta-igneous – Noril’sk).
  • Many larger IOCG deposits (meta-evaporite, brine and hydrothermal).
  • This enrichment runs counter to a proportion of 2% of the world's Phanerozoic sediments. The exact why or how of these associations is still not well understood. Most geologists working with oil, gas or metal buildups in a salt-rich basin will come to have a suspicion, and for some, the conviction, that salt or its subsurface alteration plays a role in defining the position or enrichment level of the commodity of interest. Evaporite masses in the subsurface, especially if halite-dominant, enable both physical and chemical alterations, which tend to improve economic prospectivity. The unique properties of salt in the diagenetic realm tend to facilitate, focus and stabilise processes that lead to elevated levels of accumulations of various commodities (Figure 8). That is, evaporites in a basin tend to enhance the volume of oil, gas or metals in an accumulation, but are not necessarily the direct cause of the accumulation/precipitation.

    Evaporite-hydrocarbon association

    When the reservoired hydrocarbon below a salt seal is oil, little or no leakage can take place through a laterally continuous evaporite. Even when the reservoired hydrocarbon is methane, little or no loss occurs, even by diffusion (Ehgartner et al., 1998). The much greater efficiency of evaporite seals, compared to shales, is clearly seen in the total hydrocarbon volumes held back by the two lithologies. Total worldwide shale sediment volume in the Earth's sedimentary crust is more that an order of magnitude greater than that of evaporites, yet the split between reservoired hydrocarbons below a shale or an evaporite seal is roughly 50:50 (Grunau, 1987). Likewise, typical volumes of salt-sealed giant fields dis-covered in the last decade are much grander than that held in mudrock-sealed systems (Bai and Yu, 2014).

    Of the 120 giant oil and gas fields discovered in the period 2000 - 2012 some 54.6 % are hosted in marine carbonates and 12% in lacustrine carbonates, meaning less than a third of new giant discoveries are in siliciclastic reservoirs (Figure 9a). Some 56% of these oil and gas giants have an evaporite seal, with 82% of the marine carbonates having an evaporite seal and 91% of the lacustrine carbonates having an evaporite seal (Figure 9b, c). Clearly, carbonate reservoirs with evaporite seals constitute most of the giant oil and gas discoveries across the period 2000-2012, and the proportions of this association are likely to increase in conventional discoveries across the next decades (Bai and Yu, 2014).

    Of these 120 oil and gas giants, one is a megagiant field with a recoverable reserve of 50 Bboe (Billion barrels oil equivalent) or more, and 6 are supergiant fields with recoverable reserves of 5 Bboe or more. The seven largest fields are: Galkynysh (aka Yolotan or Osman) gas field in the Ama-Darya Basin of Turkmenistan with an original 2P reserve of 67.1 Bboe, making it likely the second largest known gas field in the world; Kashagan oil field (18.1 Bboe) in the North Caspian basin; Kish 2 gas field in the Arabian Basin (Iran): Lula (previously known as Tupi), Franco and Libra oil fields in the Santos Basin of offshore Brazil; and Sulige gas field (5.7 Bboe) in the Ordos Basin, China. In this listing, the geology of Galkynysh is not yet reliably published, but of the remaining 6, only Sulige does not have an evaporite association.

    Typical bedded evaporite seals, especially beds composed of monomineralogic assemblages such as massive nodular anhydrite or massive halite and have measured entry pressures more than 3000 psi. Most impure evaporite beds have entry pressures greater than 1000 psi, as do many evaporite-plugged reflux dolomites. Contrast this with most shales, which tend to be water-bearing (mostly bound and structural water), with typical entry pressures between 900 and 1500 psi (Sneider et al., 1997). Although such shales are respectable seals, over time shale allows substantial diffu-sive leakage of methane and even liquid hydrocarbons via inherent microporosity, less so if the shales are organic-rich. Even ignoring halite's ability to reanneal and flow under stress, any evaporite seal has much lower intrinsic permeability than shale, and this helps maintain its seal integrity. With permeabilities of 10-7 md, a hydraulic gradient of 0.01, and a porosity less than  0.01, a brine would take somewhere between 3 and 30 million years to flow 1 metre into an unfractured halite seal. For anhydrite, which has permeability some 100 times higher than halite, a brine would take between 30,000 and 300,000 years to flow 1 metre into the seal (Beauheim and Roberts, 2002).

    Evaporites are excellent longterm seals to substantial hydrocarbon columns. This applies to any siliciclastic or carbonate reservoir adjacent to bedded or halokinetic salty seals (Figure 10). The exemplary ability of bedded evaporites to act as seals holding back massive hydrocarbon columns is clearly seen in the Middle East, where Ghawar, the world’s largest oil field, is sealed by bedded anhydrites of the Jurassic Arab D Formation and the overlying Hith Anhydrite seal. This evaporite seal association in Ghawar holds back an estimated remaining reserve of more than 100-200 billion barrels. Bedded platform anhydrites also seal Safaniya, the world’s largest offshore field, also in Saudi Arabia, with estimated reserves of more than 25-30 billion barrels of oil and 5 billion cubic feet of natural gas. Likewise, Permian platform anhydrites are the regional seal to North Field in offshore Qatar, the world’s largest single gas field (non-associated gas) with more than 500 tcf of reserves (Alsharhan and Nairn, 1997). This gas is reservoired and sealed in the evaporitic dolomites of the Permian Khuff Formation and recent announcements by the Qatari Government have postulated more than 900 tcf of certified non-associated gas sealed in the North Field structure.

    In terms of physical processes, buried near-pure halite beds and masses (rock salt) in the sedimentary realm are rheologically unusual compared to other nearby sediments, in that at geological time scales rock salt can set up deformation responses that mimic Newtonian fluid responses. At the same time, intercrystalline textures in the flowing salt mass maintain seal integrity, even as local crystals dissolve and reprecipitate. That is, down to depths of 6-8 km rock salt flows and maintains seal integrity, while adjacent non-salt sediments tend to fault and fracture. The ability of bedded salt to seal (Figure 10a) and of flowing salt to create and seal supra-, intra- and sub-salt reservoirs is well documented (Figure 10b).

    Dense evaporitic brines can pass through adjacent or underlying sediments both at the time the bedded salts are accumulating, or later as a subsurface salt mass dissolves. For example, chemical responses, created within the set-up hydrology of widespread salt deposition and early burial, drive brine reflux, moving magnesium-rich modified seawater brines into and through the underlying carbonate sediments. This hydrology crafts broad reflux dolomite haloes, with local burial anhydrite patches and overall improves intercrystalline connectivity, at least until the dolomite reservoir becomes overdolomitised (Figure 11). Once this happens, all effective polyhedral porosity is lost in the dolomite and the reservoir potential is destroyed. Later dissolution brines can create hydrothermal waters capable of leaching and burial dolomitization, as in the burial-dolomite reservoirs of the Western Canada Sedimentary Basin (Davies and Smith, 2006).

    Organic-hydrocarbon association

    As long ago as the middle of last century, Weeks (1961) emphasised the importance of evaporites as seals to many of the world's major hydrocarbon accumulations. He also pointed out that many of the cycles of deposition that involve organic-rich carbonate marls or muds also end with evaporites. Clearly, in evaporitic settings, there is an association with Type I-II hydrogen-prone kerogens in mesohaline source rocks, and this is related to the ability of halotolerant photosynthetic algae and cyanobacteria to flourish in periodically mesohaline waters. Such kerogens tend to be oil-prone rather than gas prone and typified by long-chain hydrocarbons (Warren, 2011).

    Much of the mesohaline organic matter preserved in evaporitic carbonates, and the resulting source rocks, originated as planktonic blooms (pelagic “rain from heaven”) or from the benthic biomass (“in situ” accumulations). Such organics typically settled out as seriate pulses of organic matter (often pelleted and laminated) that sank to the bottom of a layered brine column. Each pulse was tied to a short period when surface brines were diluted and halotolerant producers (mostly cyanobacteria and algae) flourished in the freshened lit zone. That is, laminated mesohaline mudstones that constitute most evaporitic source rocks reflect biological responses to conditions of “feast or famine” in vacillatingly-layered brine bodies (Figure 12; Warren, 1986, 2011).

    Worldwide, studies of schizohaline evaporitic basins have shown that organic-rich mesohaline sediments can accumulate beneath ephemeral surface brines in salterns, or in basin and slope settings in both marine and continental settings (Figure 12; Kirkland and Evans, 1981; Oehler, 1984; Warren, 1986, 2011; Rouchy, 1988; Busson, 1992). The most prolific accumulations of organics in ancient evaporitic settings tend to be laminated micritic carbonates deposited be-neath intermittently stratified moderately-saline (mesohaline) anoxic water columns of varying brine depth.

    There are three, possibly four, major mesohaline density-stratified settings where organic-rich laminites (source rocks) accumulate in saline environments that are also associated with, or evolve into, evaporite deposits (Warren, 2011):

    1) Basin-centre lows in marine-fed evaporitic drawdown basins (basinwide salts).

    2) Mesohaline intrashelf lows atop epeiric evaporitic platforms.

    3) Saline-bottomed lows in perennial underfilled saline lacustrine basins.

    4) Closed seafloor depressions in halokinetic deepwater marine slope and rise terrains.

    The Metal-evaporite association

    In addition to the evaporite hydrocarbon association, there is an association of evaporite settings with the larger of the known MVT, Sedex, Stratiform Sedimentary Copper and some IOCG associations (Figure 13; see Warren 2016, Chapter 15 and 16 for detailed case histories and models). The role of evaporites in focusing metalliferous ore accumulations is two-fold; 1) In solution (halite-dominant precursor) they can act as chloride-rich metal carriers and 2) Locally, as beds or masses (especially of CaSO4), their dissolution products, especially if trapped, can supply sulphur (mostly as bacteriogenic or thermogenic H2S) and also set up chemical interfaces that act as foci in the setup of brine mixing conditions suitable for precipitation of metal sulphides or native elements. Hence, most evaporite-associated ore systems tend to be epigenetic, rather than syngenetic. Subsurface salt beds and masses are merely the solid part of a large ionic recycling system; dissolved metals are another part, and zones of mixing between the two are typically sites where ores tend to accumulate. Halokinesis steadies the position of a redox interface, tied to a salt dissolution brine halo, and enables an extended phase of focus to a metal precipitative (redox) interface at a stable location in subsurface earth-space (Warren, 2016).

    At the world-scale, evaporite-associated metalliferous systems are driven by plate tectonics. Halite-dominated sequences, deposited in the drawdown basin centres, tend to dissolve in burial, and so supply chloride ions to the brine system. Salt beds that are thick enough tend to flow and so focus the upward, and centripetal passage of basinal and hydrothermal fluid flows. Dissolving gypsum or anhydrite beds, typically deposited higher on the basin platform or diagenetically accumulated along salt dissolution edges and touchdowns can supply sulphur, via bacterial or thermochemical sulphate reduction, while simultaneously focusing metalliferous brine flows into the precipitation interface.

    When the chemistries of the dissolving salt beds and the metal carriers interact so that redox fronts, salinity contrasts, and other precipitative interfaces are set up, an ore deposit can form. Thus, in base and precious metal exploration within evaporitic terranes, we are ultimately searching for those parts of a subsurface ionic cycling system where the salt dissolution, salt beds and metal systems have interacted to create economic levels of metalliferous precipitates.

    If salt is more than a seal, then......

    From a time in the 1950's and 1960's when evaporites were mostly seen as seals, knowledge systems developed over the next five decades now allow us to do much more with respect to predictive industrial geology, centred on the following evaporite facts:

    1) The distribution of evaporite depositional textures maps paleotopography across the underlying carbonates and siliciclastics. So, depositional signatures preserved in the seal can be used to map reservoir quality trends in terms of reflux dolomite intensity, anhydrite cement-patch distributions and zones of subaerial diagenesis

    2) Thick beds of halite tend to accumulate in particular plate tectonic settings. Beds deposited in plate-edge saline sumps tend to flow via sediment loading and extension, without later superimposed tectonic stresses. In contrast, bed deposited in intracratonic settings tend to require later externally-imposed tectonic stress in order to flow. Hydrocarbon trapping geometries are tied to particular styles of salt tectonics

    3) Mesohaline source rock distribution is related to evaporite basin architecture and likely fluid escape pathways, which in turn are related to seal type and timing of salt dissolution or halokinetic withdrawal.

    4) Potash ore quality controls are related to the timing of various brine crossflows generated during deposition, mesogenesis and telogenesis.

    5) Positions of likely base metal and copper accumulations relate to set-ups of dissolution-related of redox interfaces, mesogenetic cross-flows, and in some cases, halokinetic geometries. At the plate tectonic scale, these accumulations occur at particular hydrological interfaces within the basin architecture.


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    Evaporites and climate: Part 2 of 2 - Ancient evaporites

    John Warren - Saturday, February 25, 2017


    Evaporites, along with coal and bauxites, are sediments considered to be climate sensitive. Ancient evaporite distribution and associated paleolatitudes are used to reconstruct the distribution of the world's arid belts across time. As we saw in Part 1 (Salty Matters, Tuesday, January 31, 2017), thick, widespread evaporite deposit are essentially a result of the atmospheric circulation of the Hadley cells. That is, locations of subtropical dry zones and tropical/subtropical deserts of the globe are mostly determined by the positions of subsiding branches of cool, dry descending air in a Hadley cells (aka Trade Wind Belt; Lu et al., 2007; Crowley and North, 1991) within low-lying regions tied a sufficient supply of mother brine. Thus, climate plus ongoing brine supply are the underlying factors controlling locales of significant evaporite deposition (Ziegler et al., 1981). The previous article focused on regional and local climatic controls across the Quaternary (Salty Matters, Tuesday, January 31, 2017). This article extends the time frame for the evaporite/climate association across the Phanerozoic and into the Precambrian.

    As we move back in time, we move out of an icehouse-dominated world climate, with permanent ice caps waxing and waning at the world's poles, into greenhouse-dominant world climates. In greenhouse times there are no permanent polar ice sheet and glaciers occurred only in some high-altitude mountainous belts (Figure 1a, b). The transition into greenhouse climate changes the dominant 4th-order eustatic style from 100m amplitude changes every 100,000 years or so into 4th-order responses with much lower 3-5 m amplitude oscillations every 100,000 years (Figure 2a, b). Lack of polar caps raises world sealevel on the order of 40-50 metres. Thus, even without tectonic considerations, there is less continental freeboard in greenhouse times and increased the potential for significant cratonic coverage by epicontinental and pericontinental seaways (Warren, 2016; Chapter 5).

    Looking back in time, beyond the last few million years of the Quaternary, means our conceptual models must encompass a broader range of tectonic settings as well as changes in the rates of seafloor spreading, supercontinents, and times of significant igneous outpourings (superplumes). All these additional world-scale variables across a longer set of available time explain the greater range of climate potentials on the earth, compared to anything that has occurred in the time-limited base seen in icehouse-dominant Quaternary climate spreads. The last few million years of the current icehouse mode is inhabited by the human species or its primate ancestors. The previous Icehouse dominant mode was in the Carboniferous-Mid Permian when the dominant land animals were amphibians and primitive reptiles.
    One the first questions this broader ancient climate spectrum, tied to evaporites over deep time, offers up is; "How do the positions of Hadley cells vary across geological time frames?" In part 1, we saw how the rise of the Himalayas deflected a belt of cool, dry descending air much further south toward the Equator. Moving back in time creates a broader scaffolding for documenting climate variation, in part driven by the rise and fall of mountain ranges, but also influenced by increases in the rate of seafloor spreading driving ocean basin shallowing and by changes in atmospheric/seawater compositions and temperatures.

    Hadley cells and latitudinal variability over time
    According to Chen Xu et al., 2013 and Boucot et al., 2013, much of the world-scale Phanerozoic distribution of significant bedded evaporite accumulations indicates the ongoing presence of two mid-latitude arid belts, presumably situated beneath Hadley Cells. There are exceptions in locales generated in local rain shadows with orographic control provided by neighbouring mountain ranges. However, in the later Permian, through the Triassic, and much of the Jurassic the two formerly mid-latitude Hadley Cells merged over the more central Pangaeanic regions of Africa, Europe and the adjacent Americas, to form an arid belt that also encompassed low-latitude, equatorial arid regions (Figure 3). Across the continental interior of the Pangean supercontinent, this arid to hyperarid equatorial belt in the supercontinent interior prevented the formation of climate-sensitive sediments that are more typical of humid equatorial conditions that deposit coals, kaolinites, lateritic materials and bauxites. However, in this same time interval, these more humid sediment products are typically present at low latitudes of these time slices adjacent to the Panthalassic ocean.
    That is, in the absence of an equator-spanning supercontinent, low latitudes, typically imply humid and non-seasonal tropical conditions throughout much of the Phanerozoic, as we see today. But the assembly of the Pangaean supercontinent disrupted this latitudinally-zoned atmospheric circulation, replacing it with a progressively more monsoonal (seasonal) circulation and more arid, at times hyperarid, conditions in the equatorial continental interior of Pangaea (Parrish, 1993). The Pangean supercontinent reached its maximum areal extent in the Triassic and was associated with what is known as the Pangaean Megamonsoon. There were immense arid regions across the interior regions of the supercontinent that were nearly uninhabitable, with scorching days and frigid nights. However, Panthalassian coasts still experienced seasonality, transitioning from rainy weather in the summer to dry conditions during the winter and the associated accumulation of humid sediments (Figure 3b' Boucot et al., 2013). Megamonsoon aridity is evidenced not just in the accumulation of low-latitude bedded evaporite deposits. Low latitude continental aridity also drove the accumulation of thick, widespread low-latitude desert redbeds, sourced by eolian, not fluvial, detrital transport (Sweet at al., 2013) and the precipitation of bedded salt crusts in ephemeral saline lakes under exceptionally-high surface temperatures of up to 73°C (Zambito and Benison, 2013).
    Paleolatitude reconstructions Chen-Xu et al. (2013) show these continental interior arid belts in low latitude tropical-subtropical regions persisted from the Permian to the Early Cretaceous. Reunion of the humid regions from both sides of Pangaea by the early Late Cretaceous formed a through-going low latitude humid tropical-subtropical belt. This coincides with the disaggregation of Pangaean supercontinent, as the initial stages of a modern latitudinal climate belt distribution pattern emerged, tied to latitudinally-restricted evaporites that continue to the present (Figure 4).

    Tectonism and eustacy in arid climates drive the formation of mega-evaporite basins
    Within this Phanerozoic climatic framework, there are times when significant volumes of evaporites form what are know as saline giants, or megahalite/megasulphates deposits. These massive accumulations of salts formed beneath arid climates that can span both greenhouse and icehouse climates (Figure 1; Warren, 2010, 2016). Ancient marine saline giants (megahalites and megasulphates) accrued in either of two plate-scale settings, which at times merged into one another, namely; 1) Platform evaporites (Figure 5) and, 2) Basinwide evaporites (Figure 6).

    The first major contrast with nonmarine continental dominance in Quaternary evaporite settings is the fact that platform evaporites require greenhouse eustasy a and a marine feed, the second is that basinwide evaporites require tectonically- and hydrographically-isolated widespread subsealevel depressions, typically found along plate edges with continent-continent proximity in regions with a marine seepage feed and/or periodic marine overflows (Figure 6). Neither platform or basinwide conditions are present on the current earth surface. For basinwides, suitable hydrologic conditions were last present during the Messinian Salinity Crisis in the Mediterranean region, and platform evaporite settings were last present on earth across large parts of the Middle East carbonate platform during the Eocene (Tables 1, 2). There is a third group of ancient evaporite deposits; it encompasses all nonmarine lacustrine beds past and present (Table 3). This group has same-scale modern-ancient counterparts, unlike ancient marine platform and basinwide evaporites (Warren, 2010, 2016). Interestingly, the lacustrine depositional style for bedded salt accumulation dominates in the icehouse climate that is the Quaternary, and so biases a strictly uniformitarian view of the past with respect to the relative proportions of nonmarine versus marine evaporite volumes (see Part 1; Salty Matters, Tuesday, January 31, 2017).

    Platform evaporites
    Are made up of stratiform beds, usually <50 m thick and composed of stacked <1 to 5 m thick parasequences or evaporite cycles, with a variably-present restricted-marine carbonate unit at a cycle base (Table 1). Salts were deposited as mixed evaporitic mudflat and saltern evaporites, sometimes with local accumulations of bittern salts. Typically, platform salts were deposited in laterally extensive (>50-100 km wide), hydrographically-isolated, subsealevel marine-seepage lagoons (salterns) and evaporitic mudflats (sabkhas and salinas). These regions have no same-scale modern counterparts and extended as widespread depositional sheets across large portions of hydrographically isolated marine platform areas that passed seaward across a subaerial seepage barrier into open marine sediments (Figure 5). In marine margin epeiric settings, such as the Jurassic Arab/Hith and Permian Khuff cycles of the Middle East or the Cretaceous Ferry Lake Anhydrite in the Gulf of Mexico, these platform evaporites are intercalated with shoalwater marine-influenced carbonate shelf/ramp sediments, which in turn pass basinward across a subaerial sill into open marine carbonates. Landward they pass into arid zone continental siliciclastics or carbonate mudflats.


    Platform evaporite deposition occurred in both pericontinental and epicontinental settings, at times of low-amplitude 4th and 5th order sealevel changes, which typify greenhouse eustasy (Figure 5; Warren, 2010). Platform evaporites also typify the saline stages of some intracratonic basins. Platform evaporites cannot form in the high-amplitude, high-frequency sealevel changes of Icehouse eustasy. The 100m+ amplitude oscillations of Icehouse times mean sealevel falls off the shelf edge every 100,000 years, so any evaporite that had formed on the platform is subaerially exposed and leached. Fourth order high-amplitude icehouse eustatic cycles also tend to prevent laterally-continuous carbonate sediment barriers forming at the top of the shelf to slope break, and so icehouse evaporite systems tend not to be hydro-graphically isolated (drawdown) at the platform scale. Rather icehouse eustasy favours nonmarine evaporites as the dominant style, along with small ephemeral marine-margin salt bodies, as seen today in the bedded Holocene halites and gypsums of Lake Macleod in coastal West Australia (Part 1; Salty Matters, Tuesday, January 31, 2017).
    Ancient platform evaporite successions may contain halite beds, especially in intracratonic basinwide settings, but periplatform settings, outside of intracratonic basins, are typically dominated by 5–40 m thick Ca-sulphate beds intercalated with normal-marine platform carbonates (Table 1). The lateral extent of these epeiric platform sulphate bodies, like the Middle Anhydrite Member of the Permian Khuff Fm. of Saudi Arabia and the UAE, with a current area of more than 1,206,700 sq. km., constitute some of the most aerially-extensive evaporite beds ever deposited.

    Basinwide evaporites
    Are made up of thick evaporite units >50–100 m thick made up of varying combinations of deepwater and shallow water evaporites (Figure 1; Table 2). They retain textural evidence of different but synchronous local depositional settings, including mudflat, saltern, slope and basin (Figure 6). When basinwide evaporite deposition occurs, the whole basin hydrology is evaporitic, holomictic, and typically saturated with the same mineral phase across vast areas of the basin floor, as seen on a much smaller scale today in the Dead Sea basin. The Dead Sea currently has halite forming simultaneously as; 1) decimeter-thick chevron-dominated beds on the saline-pan floor of the shallow parts around the basin edge in waters typically less than 1-10 metres deep, and 2) as coarse inclusion-poor crystal meshworks of halite on the deep basin floor that sits below a halite-saturated brine column up to hundreds of metres deep. Ancient basinwide successions are usually dominated by thick massive salt beds, generally more than 100-500 m thick. Deposits are made up of stacked thick halite beds, but can also contain substantial volumes of thick-bedded Ca-sulphate and evaporitic carbonate, as in the intracratonic basinwide accumulations of the Delaware and Otto Fiord Ba-sins (Table 2).
    Owing to inherent purity and thickness of the deposited halite, many halite-dominant basinwide beds are also remobilized, via loading or tectonics, into various halokinetic geometries (Hudec and Jackson, 2007). Some basinwide systems (mostly marine-fed intracratonic settings) entrain significant accumulations of marine-fed potash salts, as in the Devonian Prairie Evaporite of western Canada. In contrast, all Quaternary examples of commercial potash deposits are accumulating in continental lacustrine systems (Warren 2016; Chapter 11).


    Basinwide evaporite deposits are the result of a combination of tectonic and hydrological circumstances that are not currently active on the world’s surface (Figure 1). They were last active in the Late Miocene (Messinian), in association with soft-suture collision basins tied to the Alpine-Himalaya orogenic belt, and in Middle Miocene (Badenian) basins developed in the early rift stages of the Red Sea. Basinwide systems will be active again in the future at sites and times of appropriate plate-plate interaction, when two continental plate edges are nearby, and the intervening seafloor is in or near a plate-edge rift or suture and is both subsealevel and hydrographically isolated (Figure 6). Unlike most platform evaporites, basinwides do not require greenhouse eustacy, only the appropriate association of arid climate and tectonics. The latter sets up a deep hydrographically-isolated subsealevel tectonic depression with a geohydrology that can draw on a huge reserve of marine mother brine in the nearby ocean. For this reason, saline giants tend to form at times of plate-scale continent-continent proximity and so occur mostly in craton-margin settings.


    Lacustrine (nonmarine) evaporites
    Quaternary continental playa/lacustrine are constructed of stratiform salt units, with the greater volume of saline sediment accumulating in lower, more-saline portions of the lacustrine landscape. Beds are usually dominated by nodular gypsum and displacive halite, deposited in extensive evaporitic mudflats and saltpans with textures heavily overprinted by capillary wicking, rather than as bedded bottom-nucleated layers on the subaqueous floors of perennial brine lakes (Ruch et al., 2012). In ancient counterparts, the total saline lacustrine thickness ranged from meters to hundreds of meters, with lateral extents measured in tens to hundreds of kilometres (Table 3). Lacustrine salt beds are separated vertically, and usually surrounded by, deposits of lacustrine muds, alluvial fans, ephemeral streams, sheet floods, eolian sands, and redbeds. As today, ancient lacustrine salts accumulated in endorheic or highly restricted discharge basins, with perennial saline water masses tending to occur in the drainage sumps of steep-sided drainage basins (Warren, 2010, 2016). Saline lake basins accumulating gypsum, or more saline salts like halite or glauberite, typically have a shallow water table in peripheral saline mudflat areas and so are dominated by continental sabkha textures. Nearby is the lowermost part of the lacustrine depression or sump where deposition is typified by ephemeral ponded brine pan deposits, rather than permanent saline waters.
    Saline lacustrine mineralogies depend on compositions of inflowing waters, so depositional sumps in regions with non-marine ionic proportions in the feeder inflow, accumulate thick sequences of nonmarine bedded salts dominated by trona, glauberite, and thenardite. In contrast, nonmarine areas with thalassic (seawater-like) inflows tend to accumulate more typical sequences of halite, gypsum, and anhydrite.
    Across the Pliocene-Quaternary icehouse, less-saline perennial saline-lake beds tend to occur during more humid climate periods in the same continental-lacustrine depressions where saline-pan beds form (e.g., Lake Magadi, Great Salt Lake, Lake Urmia). On a smaller scale, in some modern saline lake basins, parts of the lake floor can be permanently located below the water surface (Northern Basin in the Dead Sea or Lake Asal). In some modern saline sumps dominated by mudflats, a perennial saline lake water mass is located toward the edge of a more central salt-flat zone, forming a perennial water filled “moat” facies surrounding a seasonally desiccated saline pan (as in Salar, de Atacama, Salar de Uyuni, Lake Magadi, Lake Natron). These permanent to near-permanent saline water “moat” regions are typically created where fresher inflows encounter saltier beds of the lake centre, dissolve them, and so form water-filled peripheral depressions. Bottom sediment in the moats tends to be mesohaline carbonate laminites, which can contain TOC levels as high as 12%.
    High-water stage perennial saline lacustrine sediments tend to be carbonate-rich or silica-rich (diatomaceous) laminites. Ancient examples of large saline lacustrine deposits made up of alternating humid and desiccated lacustrine units include the Eocene Green River Formation of Wyoming and the Permian Pingdiquan Formation of the Junggar Basin, China (Table 4). Evaporites deposited in a suprasealevel lacustrine basin (especially Neogene deposits) have numerous same-scale Quaternary analogues, unlike the more voluminous ancient marine platform and basinwide evaporites (Figure 7). Clearly, across the Quaternary, saline continent lacustrine settings possess areas of bedded salt accumulation that are far greater than those of any contemporaneous marine-fed salt sumps (Part 1; Salty Matters, Tuesday, January 31, 2017). But in ancient climes, especially during in the continental interior of the Pangean supercontinent (mid Permian to Triassic), regions of continental interior sabkhas and saline pans had areas far greater than any seen in Quaternary continental saline sumps (Zambia and Benison, 2013).

    Is the present-day climate the key to evaporite understanding?
    This short answer is yes, Hadley Cells across the Phanerozoic are mostly tied to climate belts that maintain sub-tropical positions, but to this notion, we must add a geological context. Today we are living in an icehouse climate mode and have been for the last 12 Ma. It is tied to the presence of polar ice-sheets that wax and wane over 100,000-year time frames, so moving the position of the Hadley cells and changing the intensity of atmospheric circulation. In this icehouse climate, large eustatically-controlled marine-fed evaporite deposits are not preserved, as sea level falls off the continental shelf every 100,000 years or so. The world's largest bedded salt deposits formed sometime in the last 2 million years, are found in continental interiors, typically in endorheic tectonic sumps in either hot arid or steppe climate settings, often with salt diapirs outcropping or subcropping in the drainage basin and the basin floor can be located at elevations well above sealevel (Part 1; Salty Matters, Tuesday, January 31, 2017).
    As we move back into much of Phanerozoic time, we see world climates dominated by greenhouse modes, with shorter episodes of polar ice-sheets and icehouse climates in the Carboniferous-Early Permian ≈ 50 million years long, and the Ordovician, some 15 million years long (Figure 1). Greenhouse climate lacks permanent polar ice sheets, so sea-levels are higher, and 4th-order eustatic amplitudes in sea level are much less (a few meters versus hundred meters plus changes). Greenhouse sets up epeiric and intercontinental seaways that when hydrographically isolated, but still marine-fed, can deposit huge areas of platform evaporites centred in isolated seepage-fed sub-sealevel sumps. These platform deposits can also form outside of Greenhouse times in marine-fed tectonically-induced intracratonic sumps.
    Basinwide evaporite deposits span icehouse and greenhouse mode arid belts, whenever a marine-fed subsealevel tectonic sump forms at positions of continent-continental proximity in an arid belt. Across much of the Phanerozoic, basinwide deposits typically accumulated beneath subtropical belts of cool, dry descending air set up in a Hadley cell, and so are located north and south of a tropical equatorial belt. But the accretion of the Pangaean supercontinent (Carboniferous to Jurassic) set up conditions of continentality and orographic shadowing that allowed an arid saline belt to span the hyperarid interior of the supercontinent.

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    Evaporites and climate: Part 1 of 2 - Are modern deserts the key?

    John Warren - Tuesday, January 31, 2017

    Salt deposits and deserts

    Much of the geological literature presumes that thick sequences of bedded Phanerozoic evaporites accumulated in hot arid zones tied to the distribution of the world’s deserts beneath regions of descending air within Hadley Cells in a latitudinal belt that is typically located 15 to 45 degrees north or south of the equator (Figure 1a: Gordon, 1975). As this sinking cool air mass approaches the landsurface beneath the descending arm of a Hadley Cell it warms, and so its moisture-carrying capacity increases. The next two articles will discuss the validity of this assumption of evaporites tying to hot arid desert belts in the trade wind belts, first, by a consideration of actual Quaternary evaporite distributions as plotted in a GIS database with modern climate overlays, then in the second article via a look at ancient salt/climate distributions.

    Significant volumes of Quaternary evaporite salts are normally interpreted as being allied to the distribution of the world’s hot arid deserts (Figure 1b). In a general way this is true, but, as Warren (2010) shows, the correlation is an oversimplification. A hot arid desert does not necessarily equate to occurrences of laterally extensive bedded evaporites; there must also be a significant long-term brine inflow to the evaporative sump, incoming waters may be meteoric, marine, a hybrid and perhaps the sump is fed brines coming from dissolution of earlier formed salts in the drainage basin, including diapiric salt (Table 1; Warren, 2016).

    Actually, there are different ways of defining a desert and by implication its associated evaporites. One accepted approach is to define desert as a terrestrial area receiving less that 250 mm (10 inches) of annual precipitation. Using this definition some 26.2% of the world’s landsurface is desert (Figure 1b). But, in terms of evaporite distribution and the economics of the associated salts, this climatic generalization related to annual rainfall conceals three significant hydrological truisms. All three need to be met in order to accumulate thick sequences of bedded salts (Warren, 2010): 1) For any substantial volume of evaporite to precipitate and be preserved, there must be a sufficient volume of cations and anions in the inflow waters to allow thick sequences of salts to form; 2) The depositional setting and its climate must be located within a longer term basin hydrology that favours preservation of the bedded salt, so the accumulating salt mass can pass into the burial realm; 3) There must be a negative water balance in the basin with the potential for more water to leave the local hydrological sump than enter. When using a rainfall (precipitation) based definition of desert, the significance of these three simple hydrological axioms and the consequences, as to where bedded evaporites accumulate, is lost in the generalization that “evaporites form in the world’s deserts.”

    Continental-interior evaporites

    In an evaporite context it is better to define and plot saline hydrologies within a climatic framework where deserts are given the same hydrological consideration that is required for evaporite salts to form. That is, a desert is an area of land where annual precipitation (inflow) is less than potential evapotranspiration (outflow). This definition of a desert is the one used by Köppen (1900). With slight modification, his climatological scheme is still in widespread use to breakout the various climatic zones across the world’s landsurface (Kottek et al., 2006). Using a Köppen climate base, figure 2 plots worldwide occurrences of modern saline depositional systems with areas greater than 250 km2. Table 1 compares characteristics of some of the larger Quaternary bedded evaporite settings in marine edge and continental interiors. These regions contain evaporite salts accumulating in saline soils, sabkhas, salinas, saline lakes, playas and salt flats, with textural forms ranging from; isolated crystals and nodules in a terrigenous matrix, to salt crusts, to stacked beds of salts that can be more than 10 metres thick. The majority of the plotted saline areas are in arid regions, as defined by Köppen (Zone B), but not all such areas of widespread salts are in deserts (as defined by Köppen) and not all are in hot climates.

    The range of large (>250 km2) saline systems in the world’s arid landscape is more climatically diverse than just evaporite occurrences within a hot arid desert (BWh), although such associations do constitute some 38% of saline occurrences (Figure 3a; Warren, 2010; 2016). Cold arid deserts (BWk) host 23% of large saline occurrences, making a combined total of 61% for large evaporite accumulation (area >250 km2) occurrences in modern arid deserts (BW group), while the arid steppes (BSh and BSk) host another 22%. In total the world’s arid climatic zones host 83% of today’s larger evaporite occurrences (Figure 3b).

    This leaves another substantial, but not widely recognized, climate zone where significant volumes of Quaternary evaporites can accumulate, this is the polar tundra (ET); an environment where some 11% of large evaporite areas occur. In terms of evaporite volumes, the polar tundra (ET) is typically an arid high altitude belt, mostly in the Horse Latitude (Trade Wind) belts, and not located in polar or near-polar higher latitudes. The lakes and saline pans of the high plateaus of the Andes (Altiplano) and the Himalayas (Tibetan Plateau) typify this style of tundra (ET) evaporite. Water may be commonplace in the ET zone, but is there mostly as ice, and cryogenic salts are commonplace (see Salty Matters, Feb 24, 2015). The remaining region where significant evaporite volumes are found, some 6% of the total of large saline occurrences is a group of deposits defined by continental interior snow climates (group D), some with hot dry summers with solar evaporites alternating with dry winters favouring the possible accumulation of cryogenic salts (e.g. Great Salt Lake, USA).

    In the Northern Hemisphere the occurrence of large evaporite systems within arid deserts and steppe climates (BW and BS settings) extends much further south toward the equator and much further poleward (from 5-55°N) than the narrower range of large evaporite occurrences and associated climates in the southern hemisphere (Figure 4). This hemispheric asymmetry in evaporite occurrence is mostly a response to world-scale adiabatic effects associated with the collision of India with Eurasia and growth of the Himalayas. Today, a Cainozoic mountain range, centred on the Himalyas, diverts world-scale atmospheric air flows from the more north-south trajectory, usually associated with Hadley Cell circulation. For example, the Kunlun Mountains, first formed some 5.3 Ma, prevents moisture from the Indian Monsoon reaching much of the adjacent Tibet Plateau. Its adiabatic rain shadow creates the Taklamakan desert, the second largest active sand desert in the world (BWk).

    The uplift of the Himalayas also creates a dry easterly jet stream, moving arid cool air across the Tibet Plateau, around the northern side of the Himalayas, and then equatorward across the Arabian Peninsula toward Somalia where it descends and gains heat. That is, this stream of cool southwesterly-flowing dry air warms as it moves across the Eastern Mediterranean land areas and so heightens existing aridity. This helps create an adiabatic desert zone that today ranges across Arabia and northern Africa almost to the Equator (Figures 5).

    In the southern hemisphere, the uplift of the Andes has formed high intermontane depressions and the allied adiabatic aridity that are cooler with lower evaporation rates and higher relied in the immediate basin compared to groundwater depressions in flatter lower-elevation continental interior deserts like the Sahara. This higher stability hydrology favours salars over dry mudflats, as typified by Salar de Atacama and Salar de Uyuni. Atacama has a Quaternary saline sediment fill made up of a more than 900 m thickness of interlayered salt and clay, while Uyuni holds a more than 120 m thick interval of interbedded salt and clay infill, with areas of 3,064 km2 and 9,654 km2 and elevations of 2250 m and 3650 m, respectively (Figure 7a). These salars are the two largest known examples of Quaternary bedded halite accumulation, worldwide. Yet neither resides in hot arid desert settings (BWh); both are located in cold arid deserts (BWk) and in actively subsiding, high altitude (>2500m) intermontane (high relief) endorheic depressions.

    Worldwide, distribution of most of the larger (>250 km2) Quaternary evaporite settings located in hot arid (BWh) desert settings, tie either to; 1) endorheic river terminations along desert margins, especially if adjacent to mountain belts (e.g. the various circum Saharan chotts, playas and sabkhas adjacent to Atlas Mountains), or 2) to ancient inherited paleodrainage depressions (as in the majority of the interior salt lakes of Australia or the southern Africa pans). Another hot arid desert (BWh) evaporite association is defined by termination outflow rims of deep artesian systems, as in Lake Eyre North, Australia (8,528 km2, with an ephemeral halite crust up to 2 m thick in its southern portion). Similar, deeply-circulating, meteoric artesian hydrologies help explain the distribution of chotts in the BWh zone of NE Africa. Unlike Atacama and Uyuni in the Andes, none of these modern BWh artesian systems preserve stacked decametre-thick salt beds, nor did they do so at any time in the Quaternary. Rather, the most extensive style of BWh salt in meteoric-fed artesian outflow zones is as dispersed crystals of gypsum and halite in a terrigenous redbed matrix (sabkha) or as visually impressive large ephemeral saline flats and pans covered by metre-scale salt crusts that dissolve and reform with the occasional decadal freshwater flood (Warren 2016; Chapter 3). Sediments of such continental groundwater outflow zones are typically reworked by eolian processes and, due to a lack of long term watertable stability, the longterm sediment fill is matrix-rich and evaporite-poor, with the Quaternary sediment column typified by episodes of deflation, driven by 10,000-100,000 year cycles of glacial-interglacial climate changes.

    It seems that to form and preserve laterally-extensive decametre-thick stacked beds of halite in a Quaternary time-framework requires an actively subsiding tectonic depression in a cooler high-altitude continental desert, where temperatures and evaporation rates are somewhat lower than in BWh settings, allowing brine to pond and remain at or near the surface for longer periods (Figures 6, 7a). But perhaps more importantly, all of the larger regions of the Quaternary world, where thick stacked bedded (not dispersed) evaporites are accumulating, are located in continental regions with drainage hinterlands where dissolution of older halokinetic marine-fed salt masses are actively supplying substantial volumes of brine to the near surface hydrology. This halokinetic-supplied set of deposits includes, Salar de Uyuni and Salar de Atacama in the Andean Altiplano, the Kavir salt lakes of Iran, the Qaidam depression of China and the Dead Sea (Table 1).

    Marine-edge evaporites

    Thick sequences of stacked Quaternary evaporite beds with a marine-brine feed are far less common and far smaller than meteoric/halokinetic Quaternary continental evaporite occurrences. Relatively few marine-fed evaporite regions exist today with areas in excess of 250 km2 (Table 1; Figures 6, 7b). By definition, in order to be able to draw on significant volumes of seawater, these basins must operate with a subsealevel hydrology. This allows large volumes of seawater to seep into the depression and evaporate. It also means most of these deposits are located near the continental edge where a freestanding mass of seawater is not too distant

    The largest known deposit of this group is Lake Macleod on the west coast of Australia, with an area of 2,067 km2 and containing a 10m-thick Holocene gypsum/halite bed (Figure 7b). It hosts a saltworks producing some 1,500,000 tonnes/year of halite from lake brines in a BWh setting (Warren, 2016). A smaller marine seepage example, with a similar Quaternary coastal carbonate dune-hosted seepage hydrology, is Lake MacDonnell near the head of the Great Australia Bight. It has an area of 451 km2, a 10m-thick fill of Holocene bedded gypsum and is located in a milder BSk setting, compared to Lake Macleod. Even so, annually, the Lake MacDonnell operation is quarrying more than 1.4 million tonnes of Holocene coarsely-crystalline near-pure gypsum and producing more that 35,000 tonnes of salt via by pan evaporation of lake brines.

    Interestingly, when the climatic settings of Holocene coastal salinas of southern and western Australia are compared, all show similar interdunal sump seepage hydrologies with unconfined calcarenite aquifers, yet it is clear that gypsum evaporite beds dominate in BSk and lower precipitation levels, with more carbonate, in Csb coastal settings typified by hot dry summers. Halite dominates the marine-fed bedded fills in BWh coastal settings, while Coorong-style meteoric-fed carbonates dominate in similar interdunal coastal seepage depressions in the more the humid and somewhat cooler Csb settings of the Coorong region.

    One of the most visually impressive marine seep systems in the world is Lake Asal, immediately inland of the coast of Djibouti, with an area of only 54 km2 it is much smaller than the 250 km2 cutoff used for this discussion. It is located in a BWh climate, similar to that of the Danakil depression, contains subaqueous textured gypsum and pan halites, and lies at the bottom of a hydrographically-isolated basalt-floored depression with a brine lake surface some 115m below sealevel (Figure 8; see Warren, 2016 Chapter 4 for geological detail). This difference in elevation between the nearby Red Sea and the lake floor drives a marine seep hydrology, so that seawater-derived groundwater escapes as springs along the basaltic lake margin. Most of the cations and anions in the seawater feed, that ultimately accumulate as salts in the subsealevel Asal sump, move lakeward via gravitational seepage along fractures in a basaltic ridge aquifer separating the Red Sea from the brine lake, in what is an actively extending continental rift that will shortly become and arm of the Red Sea.

    That there are very few large marine-fed bedded Quaternary evaporite systems is illustrated by summing the total area of large active continental-fed evaporite areas (>250km2) listed in Figure 2, which gives a total surface area worldwide in excess of 360,000 km2. In contrast, the total area of large modern marine-fed bedded salt systems, at slightly less 10,000 km2, is more than an order of magnitude smaller. This low value for large coastal marine-fed evaporite occurrences is in part because many classic coastal sabkhas, with characteristic dispersed evaporites in their supratidal sediments and long considered to be archetypal marine-fed groundwater evaporite system, are now seen as mostly continental brine-fed hydrologies.

    For example, the modern Abu Dhabi sabkha system (1,658 km2; BWh) has been shown by Wood (2010) to be a continental groundwater outflow area, where most of ions precipitating as salts in the supratidal zone are supplied by upwelling of deeply-circulated meteoric waters, not seawater flooding (Warren, 2016; Chapter 3). Similar continental hydrologies, resurging into the coastal zone. supply much of the salt accumulating in the suprasealevel sabkhas near Khobar in Saudi Arabia (BWh), Rann of Kutch in India (28,000 km2, BWh), Sabkha Matti (2,955km2, BWh) in the Emirates, and Sabkha de Ndrhamcha (634 km2, BWh) in Mauritania. And yet today all of these large sabkha occurrences are located just inland of modern coastal zones and so, without hydrological knowledge, are easily interpreted as marine. The distinction between a water budget and a salt supply budget emphasizes a general observation in Holocene evaporite systems that, in order to define the main fluid carrier for the salt volume found in any supratidal part of an evaporitic depression, it is important to understand and quantify the nature of the groundwater feed, be it by marine or nonmarine, in both continental-interior and marine-margin settings. It also means that surface runoff and seawater washovers, while at times visually impressive, typically do not supply the greater volume of preserved salts in most modern coastal sabkha systems.

    Due to the highly impervious nature of muddy sabkha sediment, an occasional seawater storm flood into a coastal margin mudflat does not mean the pooled seawater ever penetrates the underlying sabkha. Most of its solute load is deposited as an ephemeral surface crust of halite that is a few centimetres thick atop the sabkha muds (Wood et al., 2005). Based on Wood’s and other hydrological studies of modern coastal sabkhas in BWh settings, centred on the Middle East, it seems that the salt supply to most modern coastal sabkha depressions is still not at hydrological equilibrium with the present sealevel, which was reached at the beginning of the Holocene some 6,000-8,000 years ago. That is, modern zones of continental groundwater outflow along an arid coast can have watertables that are up to a metre of two above sealevel and are strongly influenced by the resurgence of seaward-flowing deeply circulating, continental groundwaters (it is axiomatic that surface runoff and local meteoric recharge tend to be limited in BWh settings). In this situation resurging waters in the sabkha tend to have a dominantly continental ionic supply for their preserved salts.

    Lastly, in terms of the climatic and hydrological pre-requisites for the accumulations of thick bedded Holocene salt deposits, one must recognize that marine-margin sabkha mudflats do not have the same geohydrology as a near-coastal, hydrographically-isolated seepage-fed sub-sealevel salina depression. A BWh or BSk near-coastal salina is slowly but continually resupplied ions via springs fed by the effusion of seawater, moving under a gravity drive (evaporative drawdown), through an aquifer that is the barrier separating the salina depression from the adjacent ocean. Until the depositional surface reaches hydrological equilibrium, it draws continually on the near limitless supply of ions in the adjacent salty reservoir that is the world’s ocean. Salts growing displacively in the supratidal parts of a sabkha cannot call on a drawdown hydrology and can only be supplied marine salts from salt spray as it is blown inland or from waters washed over the sabkha by the occasional storm driven wash-over and breakout (Figure 9; Warren and Kendall, 1985).

    So what?

    So, if we based our ideas of where ancient bedded evaporites formed on a strictly uniformitarian approach using Quaternary analogues, then the study of where bedded salts have formed best in the world over the last two million years leads to a conclusion that large bedded salt deposits are not marine fed. Rather, in the Quaternary, thick bedded salts form best in tectonically active, subsiding hypersaline groundwater sumps located in high-altitude high-relief cold-arid deserts. These depressions are hydrographically closed, with the purer, most voluminous, examples of 100m-thick evaporite successions located in tectonically active piggy-back depressions, with lake floors and hydrologies that are more than 2000 m above sea level. Their hydrologies are endorheic and strongly influenced by deeply circulating meteoric waters. In addition, the better examples of thick stacked bedded salts are supplied ions via the dissolution of older marine halokinetic salts in the surrounding drainage basin (Table 1).

    Yet, anyone working in ancient (pre-Quaternary) evaporites knows from simple arguments of salt volumetrics versus brine sources, and the nature of the enclosing sediments, that they are part of a dominantly marine basin fill. This will be the focus


    Gordon, W. A., 1975, Distribution by latitude of Phanerozoic evaporite deposits: Journal of Geology, v. 83, p. 671-684.

    Köppen, W., 1900, Versuch einer Klassification der Klimate, vorzsugsweise nach ihren Beziehungen zur Pflanzenwelt: Geographraphische Zeitschrift, v. 6, p. 593-611.

    Kottek, M., J. Grieser, C. Beck, B. Rudolf, and F. Rubel, 2006, World Map of the Köppen-Geiger climate classification updated: Meteorologische Zeitschrift, v. 15, p. 259-263.

    Stieljes, L., 1973, Evolution tectonique récente du rift d'Asal: Review Geographie Physical Geologie Dynamique, v. 15, p. 425-436.

    Warren, J. K., 2010, Evaporitic source rocks: mesohaline responses to cycles of “famine or feast” in layered brines, Quaternary carbonate and evaporite sedimentary facies and their ancient analogues, John Wiley & Sons Ltd., p. 315-392.

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

    Warren, J. K., and C. G. S. C. Kendall, 1985, Comparison of sequences formed in marine sabkha (subaerial) and salina (subaqueous) settings; modern and ancient: Bulletin American Association of Petroleum Geologists, v. 69, p. 1013-1023.

    Wood, W. W., 2010, An historical odyssey: the origin of solutes in the coastal sabkha of Abu Dhabi, United Arab Emirates, Quaternary carbonate and evaporite sedimentary facies and their ancient analogues, John Wiley & Sons Ltd., p. 243-254.

    Wood, W. W., W. E. Sanford, and S. K. Frape, 2005, Chemical openness and potential for misinterpretation of the solute environment of coastal sabkhat: Chemical Geology, v. 215, p. 361-372.

    [1] The word desert comes from the Latin dēsertum (originally "an abandoned place"), a participle of dēserere, "to abandon."


    Gases in Evaporites Part 3 of 3; Where do gases generate and reside at the scale of a salt mass or salt bed

    John Warren - Saturday, December 31, 2016

    So far we have looked at gas distribution and origins in evaporites at micro and mesoscales and have now developed sufficient understanding to extrapolate to the broader scale of architecture for a large body of salt in an evaporite. We shall do this in a classification framework of extrasalt versus diagenetic periphery versus intrasalt gas in a halokinetic salt mass (Figure 1).

    Extrasalt gas and brine intersections

    This type of gas intersection is perhaps the most damaging to a salt mine operation and tends to occur when a gas release is encountered in an expanding mining operation, or a drill hole, that lies near the salt body edge and intersects nonsalt sediments. Extrasalt fluids can be either normally pressured or overpressured depending on the connectivity of the plumbing in the extrasalt reservoir. Salt because of its excellent seal potential tends not to leak or leak only slowly, so facilitating significant pressure buildup (Warren, in press)

    The gas inflow from this type of extrasalt breach in a salt mine is typically accompanied, or followed by, a brine release that sometimes cannot be plugged, even by a combination of grouting and brine pumping. Brine inflow rates in this scenario tend to increase with time as ongoing salt dissolution is via ongoing undersaturated water crossflows and the mine or the shaft is ultimately lost to uncontrollable flooding of gas blowouts in an oil well with poor pressure control infrastructure and planning. This type of edge intersection is why a number of early attempts to construct shafts for potash mines in western Canada failed in the middle of last century. It is why freeze curtains are considered the best way to contract a shaft for a potash mine. Examples of this type of gas/brine intersection are usually tied to telogenetic fluid entry from substantial aquifer reservoirs outside the main salt mass and are discussed in detail in Warren, (2016, Chapter 13) and as a type of salt anomaly association discussed in Warren (in press).

    The extrasalt source and potential inflow volume of this form of gas (mostly methane and co-associated brine) is largely tied to maturity of hydrocarbon source rocks located external to the salt mass in both suprasalt and subsalt positions (Figure 1). In the past, unexpected extrasalt intersections of pressurised gas reservoirs during oil well drilling lead to spectacular blowouts or “gushers”, especially in situations where the salt held back a significant volume of fluid held in open fractures beneath or adjacent to a salt seal (Table 1). The fluid-focusing effects of suprasalt dome drape and associated extensional falling and gas leakage also mean “gas clouds” are common above salt domes (Warren, 2016, in press). Low σhmin leads to upward gas migration through fracturing (Dusseault et al., 2004). So, in the supradome extrasalt position, simultaneous blowout and lost circulation conditions can be encountered, as well as the problem of severely gas-cut drilling fluids. The volumes of gassy liquids held in pressurised extrasalt reservoirs can be substantial so blowouts or “gushers” can be difficult to control, as was the case with the world-famous subsalt Qom (1956) and suprasalt Macondo (2010) blowouts (Table 1). Methane and gassy liquids generated by organic maturation tend to be the dominant gases found in this situation.


    Caprock and other salt periphery-held gases

    This style of gas occurrence is in part related to gases sourced in maturing extra-salt sediments but also taps gases that are the result of the diagenetic processes that create caprocks. Caprocks are alteration and dissolution haloes to both bedded and halokinetic salt masses and so are distinct gas reservoirs compared to extrasalt sediments (Warren, 2016; Chapter 7). They are compilations of fractionated insolubles left behind at the salt dissolution interface as the edge of halite mass liquefies. Accordingly, caprocks are zoned mineralogically according rates of undersaturated fluid crossflow and in part responding to variable rates of salt rise and resupply. Anhydrite (once suspended in the mother salt) accretes at the dissolution front. Ongoing undersaturated crossflow at the outer contact of the anhydrite residue carapace alters anhydrite to calcite via bacterially- or thermochemically-driven sulphate reduction, with hydrogen sulphide as a by-product. Additional sulphate reduction can occur in the extrasalt sediment both at or near the caprock site, but also deeper or more distal positions in the extrasalt, so sulphate reduction can be a major source of the H2S gas found in the salt periphery. H2S can also migrate in a c from sulphate reduction in maturing sediments located some depth below the salt.

    Dissolution that facilitates caprock also drives the creation of vugs and fractures in the caprock, and is one of the primary controls on reservoir poroperm levels in various caprock oil and gas reservoirs discovered in the 1920s in the US Gulf Coast. Methanogenic biodegradation of the same hydrocarbons, which facilitate sulphate reduction, can generate CO2 in the caprock and extrasalt sediments (Clayton et al., 1997)

    Many salt mine problems in Germany in the early days of shaft sinking for salt mining were related to unexpected shallow gas outflows confronted within caprock-hosted gas-filled vugs and fractures encountered by the mine shaft on the way to a potash ore target (Gropp, 1919; Löffler, 1962; Baar, 1977). Likewise, the highly unpredictable distribution of gases in the shallow caprocks and salt peripheries of the US Gulf Coast were the cause of some spectacular blowouts such as Spindletop (1901) (Table 1). Because the volume of held liquids is more limited in the vugs and fractures in a caprock compared to fractured subsalt reservoirs, the rate of fluid escape in a “caprock-fed” gusher tends to lessen and even self-bridge more rapidly than when salt is sealing a fractured overpressured subsalt reservoir (days or weeks versus months). As such these intersections, if isolated from extrasalt reservoirs as not such a problem in the drilling of oil wells. In simpler, less environmentally conscious, early days of oilwell drilling in East Texas in the 1920s, “gushers” were often celebrated, tourist spots and considered a sign of the potential wealth coming to the country being drilled.

    Intrasalt gas

    This type of accumulation/intersection is often described as an intrasalt gas pocket and is typified by a high rate of gas release, that in a mine is accompanied by a rockburst, followed by a waning flow that soon reaches negligible levels as the pocket drains (see article 1 in this series). Intrasalt gas pockets can create dangerous conditions underground and lives can be lost, but in many cases after the initial blowout and subsequent stabilisation, the mine operations or oil-well drilling can continue. Gas constituents and relative proportions are more variable in intrasalt gas pockets compared to gases held in the extrasalt and the periphery. Extra-salt gases are typically dominated by methane with lesser H2S and CO2, periphery gases by H2S and methane, while intrasalt gases can be dominated by varying proportions of nitrogen, hydrogen or CO2. Methane can be a significant component in some intrasalt gas pockets, but these occurrences are usually located in salt anomalies or fractures that are in current or former connection with the salt periphery.

    Gas types and sources at the local and basin scale

    The type of gas held within and about a salt mass in a sedimentary basin is broadly related to position in the mass and proximity to a mature source rock. Herein is the problem, most of the gases that occur in various salt-mass related positions (intrasalt, extrasalt and periphery) can have multiple origins and hence multiple sources.

    Accumulations of gas with more than 95 vol.% N2 are found in most ancient salt basins and the great majority of these accumulations are hosted in intersalt and subsalt beds, with the gas occurring in both dispersed and free gas forms in the salt, as in many Zechstein potash mines of Germany and the Krasnoslobodsky Mine in the Soligorsk mining region of Russia (Tikhomirov, 2014). Nitrogen gas today constitutes around 80% of earth atmosphere where it can result from the decay of N-bearing organic matter (proteins). Ultimately, nitrogen speciates from aqueous mantle fluids in oxidised mantle wedge conditions in zones of subduction and in terms of dominance in planetary atmospheres it indicates active plate tectonics (Mikhail and Sverjensky, 2014). Nitrogen in the subsurface is large unreactive compared to oxygen and so tens to stay in its gaseous form while oxygen tens to combine into a variety of minerals. When held in a salt bed, nitrogen can be captured from the atmosphere during primary halite precipitation and stored in solution in a brine inclusion so creating a dispersed form of pressurised nitrogen. When buried salt recrystallizes during halokinesis, with flow driven by via pressure solution, inclusion contents can migrate to intercrystalline positions and from there into fractures to become free gas in the salt.

    Methane gas captured in and around a salt mass as both dispersed and be gas typically mostly comes from organic maturation. The maturing organic matter can be dispersed in the salt during primary halite precipitation, it can be held in intersalt source beds (as in the Ara Salt of Oman), or it can migrate laterally to the salt edge, along with gases and fluids rising from more deeply buried sources. Thus, the presence of oil, solid bitumen and brine inclusions, with high contents of methane in halite, does not unequivocally point to the presence of oil or gas in the underlying strata, it can be locally sourced from intersalt beds as in the Ara Salt. However, a geochemical aureole can be said to occur if hydrocarbons in the halite-hosted inclusions can genetically be linked with reservoired oil or gas. The presence of methane in salt anomalies in Louann Salt mines in the US Gulf Coast and some mines in Germany is likely related to organic maturation of deeply buried extrasalt source rocks with subsequent entrapment during halokinesis and enclosure of allochthon-suture sediments.

    Hydrogen sulphide gas (H2S) is a commonplace free gas component in regions of bacterial and thermogenic sulphate reduction. Like methane, much of its genesis is tied to organic maturation products (and sulphate reduction processes), and like methane, it can be held in salt seal traps, or in peripheral salt regions, or in intrasalt and intersalt positions and like metyhane if it escapes and ponds in an air space its release can be deadly (Table 1; Luojiazhai gas field, China). Because both bacterial and thermochemical sulphate reduction requires organic material or methane, there is a common co-occurrence of the two gases. Caprock calcite phases are largely a by-product of bacterial sulphate reduction, so there is an additional association of H2S with caprock-held occurrences. This form of H2S, along with CO2, created many problems in the early days of shaft sinking in German salt mines. More deeply sourced H2S tend to be a production of thermochemical sulphate reduction in regions where pore fluid temperatures are more than 110°C.

    Detailed study of CO2 and its associated geochemical/mineralogic haloes shows much of the CO2 held in Zechstein strata of Germany has two main sources; 1) Organic maturation and 2) carbonate rock breakdown especially in magmatic hydrothermal settings (Fischer et al., 2006). The organic-derived CO2 endmember source (with δ13C near -20‰) is present in relatively low concentrations, whereas large CO2 concentrations are derived from an endmember source with an isotope value near 0‰. Although the latter source is not unequivocally defined by its isotopic signature, such “heavy” CO2 sources are most likely attributed to heating-related carbonate decomposition processes. This, for example, explains the CO2-enriched nature of salt mines in parts if the former East Germany where Eocene intrusives are commonplace (Shofield et al., 2014).

    Hydrogen (H2) gas distribution as a major component varies across salt basins and is especially obvious in basins with significant levels of carnallite and other hydrated potassic salts. This association leads to elevated radiogenic contents tied to potassic salt units, with hydrogen gas likely derived from the radiogenic decomposition of water (see article 2 in this series). The water molecules can reside in hydrated salts or in brine inclusions in salt crystals.


    Various proportions of gases (N2, CH4, CO2, H2S, H2) held in salt as dispersed and free gas occur in all salt basins. But at the broad scale, certain gases are more common in particular basin and tectonic positions. Methane is typically enriched in parts of a basin with mature source rocks, but can also have a biogenic source. Likewise, H2S is tied to zones of organic breakdown, especially in zones of either bacterial or thermochemical sulphate reduction. CO2 can occur in salt in regions of organic degradation, but is most typical those of parts of a salt basin where igneous processes have driven to thermal and metamorphic decomposition of underlying carbonates (including marbles). Nitrogen because of its inert nature is a commonplace intrasalt gas and comes typically from zones of organic decomposition with dispersed nitrogen becoming free gas with subsequent halokinetic recrystallisation. Ongoing salt flow can drive the distribution of all dispersed salt stored gases into free gas (gas pocket) positions.


    Baar, C. A., 1977, Applied salt-rock mechanics; 1, The in-situ behavior of salt rocks: Developments in geotechnical engineering. 16a.

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    Warren, J. K., in press, Salt usually seals, but sometimes leaks: Implications for mine and cavern stability in the short and long term: Earth-Science Reviews.



    Gases in Evaporites, Part 2 of 3: Nature, distribution and sources

    John Warren - Wednesday, November 30, 2016

    This, the second of three articles on gases held within salt deposits, focuses on the types of gases found in salt and their origins. The first article (Salty Matters October 31, 2016) dealt with the impacts of intersecting gassy salt pockets during mining or drilling operations. The third will discuss the distribution of the various gases with respect to broad patterns of salt mass shape and structure (bedded, halokinetic and fractured)

    What’s the gas?

    Gases held in evaporites are typically mixtures of varying proportions of nitrogen, methane, carbon dioxide, hydrogen, hydrogen sulphide, as well as brines and minor amounts of other gases such as argon and various short chain hydrocarbons (Table 2). There is no single dominant gas stored in salt across all evaporite deposits, although a particular gas type may dominate or be more common in a particular region. For example, CO2 is commonplace in the Zechstein salts of the Wessen region of Germany (Knipping, 1989), methane is common in a number of salt dome mines in central Germany and the Five-Islands region in Louisiana, USA (Kupfer, 1990), nitrogen is dominant in other salt mines in Germany and New Mexico, while hydrogen can occur in elevated proportions in the Verkhnekamskoe salt deposits of the Ural foredeep (Savchenko, 1958).

    Before considering the distribution of the various gases, we should note that older and younger sets of gas analyses conducted over the years in various salt deposits are not necessarily directly comparable. Raman micro-spectroscopy is a modern, non-destructive method for investigating the unique content of a single inclusion in a salt crystal. There is a significant difference in terms of what is measured in analysing gas content seeping from a fissure in a salt mass or if comparisons are made with conventional wet-chemical methods which were the pre Raman-microscopy method that is sometimes still used. Wet chemical methods require sample destruction, via crushing and subsequent dissolution, prior to analysis. This can lead to the escape of a variable proportion of the volatile compounds during the crush stage, such as methane, hydrogen, ethane and aromatic hydrocarbons, especially of those components held in fissures and more open intercrystalline positions. Any wet chemical technique gives values that represent the average of all the inclusion residues and intercrystalline gases left in the studied sample, post preparation. In contrast, Raman Microspectroscopy indicates content and proportion within a single inclusion in a salt crystal. So, free gas results and wet chemical compositions, when compared to Raman microscopy determinations from inclusions, are not necessarily directly comparable. With this limitation in mind, let us now look at major gas phases occluded in salt.


    Gassy accumulations in salt with elevated levels of N2 occur in many salt basins in regions not influenced by magmatic intrusions (Table 1). In an interesting study of spectroscopic gases held in inclusions in the Zechstein salt of Germany, Siemann and Elendorff (2001) document a bipartite distribution of inclusion gases. With rare exceptions, the first group, made up of N2 and N2-O2 inclusions reveals N2/O2 ratios close to that of modern atmosphere, which they interpret as indicating trapped paleoatmosphere (Figure 1). Similar conclusions are reached in earlier studies of nitrogen gas held in Zechstein salts, using wet chemical techniques (Freyer and Wagener, 1975). The second group documented by Sieman and Elendoorff (2001) is represented by inclusions that contain mixtures of N2, CH4 occasionally H2 or H2S. The most abundant subgroups in this second group are N2-CH4 and N2-CH4-H2 mixtures, that is, the methane association (Figure 1). Siemann and Elendorff (2001) argue that these methanogenic and hydrogenic gas mixtures of the second group are the product of decomposition of organic material under anoxic subsurface conditions. They note that the methane and hydrogenic compounds, as well as some portion of the nitrogen, are not necessarily derived from decomposing organics held within the salt. They could have been generated by degassing of underlying Early Permian (Rotliegendes) or Carboniferous organic-rich sedimentary rocks with subsequent entrapment during early stages of fluid migration, possibly driven by Zechstein halokinesis.

    Different origins and timings of both main nitrogen gas groupings in inclusions in the salt host is supported by stratigraphic correlations (Siemann and Elendorff, 2001). In the stratigraphic layers which contain mainly mixtures of N2 and O2 or pure N2, inclusions of the N2-CH4-H2-H2S-group are rare (A in Figure 1) and vice versa: layers which are rich in N2-CH4-H2-H2S do not contain many pure N2-O2inclusions (B in Figure 1). The majority of layers investigated in the salt mostly contain inclusions of the N2-O2 group, sans methane. Only two anhydrite-rich layers of Zechstein 3 (Main Anhydrite and Anhydrite-intercalated Salt) contain mainly inclusions of the second group (i.e. with abundant methane) as seen in B in Figure 1. The Zechstein 3 potash seams, as well as secondary halites, contain more or less the same population of inclusions from every main group (C in Figure 1). A comparison of the gas-rich inclusions and the gases in the brine-rich inclusions of the Zechstein 2 layer, Main Rock Salt 3, also shows distinct differences. Whereas, the gas-rich inclusions are mostly of the N2-O2 grouping, the gases from the brine-rich inclusions are mostly of the N2-CH4 group, emphasizing different origins for the gas-rich and brine-rich inclusions Siemann and Elendorff (2001) conclude that the latter gas group is a product of thermally evolved anhydrite-rich parts or potash seams that have generated hydrocarbons catagenically, with these products migrating into the overlying and deforming Main Rock Salt 3.


    Work on the free gas released during mining of the Permian Starobinsky potash salt deposit in the Krasnoslobodsky Mine, Soligorsk mining region, Russia shows that the dominant free gas is nitrogen, along with a range of hydrocarbons, including methane (Figure 2; Andreyko et al., 2013). The compositional plot is based on free gases released from the main pay horizon of the Krasnoslobodsky Mine, which it the Potash Salt Horizon 3. The exploited stratigraphy is 16 to 18 m thick in the centre of the minefield and thins to 1 m thick at the edges of the ore deposit. Depth to the potash horizon varies from 477 to 848 m below the landsurface. It consists of three units: 1) top sylvinite unit, which is classified as non-commercial due to high insoluble residue content; 2) mid clay–carnallite unit, which is composed of alternating rock salt, clay and carnallite; and 3) bottom sylvinite unit, which is the main ore target and is composed of six sylvinite layers (I-VI), alternating with rock salt bands (Figure 3). The distribution of gas across the stratigraphy of units I-VI shows that the free gas yields are consistently higher in the sylvinite bands (Figure 3).


    Oxygen levels in salt are not studied in as much detail as the other gas phases due to their more benign nature when released in the subsurface. Work by Freyer and Wagener (1975) focusing on the relative proportions of oxygen to nitrogen held in Zechstein salts was consistent with the inclusions retaining the same relative proportions of the two gases as were present in the Permian atmosphere when the salts first precipitated.

    As well as being held within the salt mass, substantial nitrogen accumulations can be hosted in inter-salt and sub-salt lithologies. For example, the resources of nitrogen in the Nesson anticline in the Williston Basin are ≈53 billion m3 and held in sandstones intercalated with anhydrite in the Permian Minnelusa Fm (Marchant, 1966; Anderson and Eastwood, 1968) and those in Udmurtia in the Volga–Ural Basin are ≈33 billion m3 (Tikhomirov, 2014). In both these non-salt enclosed cases the evolution of the nitrogen gas is related to the catagenic and diagenetic evolution of organic matter. Tikhomirov (2014) concludes that nitrogen in the various subsalt fluids in the Volga–Ural Basin originates from two major sources. Most of the nitrogen in the subsalt has δ15N > 0‰ and is genetically related to concentrated calcium chloride brines, heavy oils, and bitumen in the platform portion of the basin and so ties to a catagenic origin. The other N2 source is seen in subordinate amounts of nitrogen across the basin with δ15N values < 0‰. According to Tikhomirov (2014), this second group seems to be genetically related to methane derived at significant depths in the basement lithologies of Ural Foredeep and Caspian depression (possibly a form of mantle gas?).


    Unexpected intersections with gas pockets containing significant proportions of methane can be dangerous, as evidence by the Belle Isle Salt Mine disaster in 1979 as well as others (see article 1). Many methane (earth-damp) intersections and rockbursts in US Gulf Coast salt mines can be tied to proximity to a shaley salt anomaly (Molinda 1988; Kupfer 1990).

    Methane contents of normal salt (non-anomaly salt) in salt domes of the Five-Islands region of the US Gulf Coast were typically low (Kupfer, 1990). For example, the majority of the samples of normal salt, as tested by Schatzel and Hyman, (1984), contained less than 0.01 cm3 methane per 100 g NaCI. Although there can be wide ranges of methane enrichment in normal versus outburst salts, outburst salts are typified by increases in halite crystal size, the number of included methane gas bubbles, contorted cleavage surfaces related to increased overpressured gas contents, and an increase in clay impurities in some of the more methane-rich salt samples.


    Probably the most detailed study of controls on methane distribution in domal salt was conducted at the Cote Blanche salt mine in southern Louisiana (Molinda, 1988). Because outbursts were the primary mode of methane emission into the mine, he mapped more than 80 outbursts, ranging in size from 1 to 50 ft in diameter. The outbursts were aligned and elongate parallel to the direction of salt layering and such zones correlate well with high methane content (Figure 4). Halite crystal size abruptly increased upon entry into gassy zones subject to rockburst. The intensity of folding and kinking of the salt layering within the outburst zone also increased. The interlayered sand, shown in Figure 4, also occurred throughout the mine and not just in the mapped area shown, but was not a significant source of methane. Molinda (1988) and Schatzel and Hyman (188) all concluded that not all rockbursts were hosted by coarsely crystalline fine-grained salt, so the absence of coarsely crystalline salt may not be an indication that a rockburst cannot occur, although it is less likely. Sampling the salt for methane levels may be a better approach for rockburst prediction.

    In some methane occurrences in Europe (in addition to generation from clayey intrasalt organic entraining bands) there is a further association with igneous-driven volatilization from nearby, typically underlying, coaly deposits. This igneous association with coals and carbonates likely creates an additional association with CO2 and possibly H2S.


    Many CO2 rich gas intersections tie to regions that have been heated or cross-cut with igneous intrusives. For example, many of the CO2-bearing gas mixtures that were encountered in the Werra region during the initial exploratory drillings for potash salts(Table 1 in article 1; Frantzen, 1894). In 1901, shortly after mining at Hämbach had begun, coincident intersections of basalt dykes and releases of gas were observed (Gropp, 1919). Dietz (1928a,b) noted that a fluid phase was always involved in the fixation of the CO2gas mixtures in the Zechstein evaporites, while Bessert (1933) reported on the enrichment of anhydrite, kainite, and polyhalite at the contact with the basaltic intrusive. Accumulations of CO2-rich free gas in many Wessen mines became a safety issue and many subsequent studies underlined the association of CO2 enriched gases with basalt occurrences (Knipping, 1989). In almost all instances in the Zechstein where native sulphur forms the at the contact of a basaltic dyke, knistersalz dominates the evaporite portion of the samples. According to Ackermann et al. (1964) gas-bearing drill core samples collected in the Zechstein K1Th unit (carnallitite, sylvinite) in the Marx-Engels mine (formerly Menzengraben, East Germany) contained up to 0.6 - 14.0 ml gas/100 gm rock, with an average of 3.6 ml of gas fixed in 100 g of salt rock (Table 1)of. On average, the gas inclusions were composed of 84 vol% CO2. Knipping (1989) concludes that quantities of volatile phases (mainly H20 and CO2) penetrated the evaporites during intrusion of basaltic melts. These gases influenced mineral reactions, particularly when intersecting with reactive K-Mg rock layers of the Hessen (K1H) and Thuringen (K1Th) potash seams in the former East Germany. The intensity of this reaction was likely greater when the evaporite layers contain hydrated salts such as carnallite and kainite. Such salts tend to release large volumes of water at relatively low temperatures when heat by a nearby intrusive (Warren, 2016; Chapter 16; Schofield et al., 2014). In doing so, significant volumes of CO2 enriched gases were trapped in the altered and recrystallising evaporites, so forming knistersalz.

    While discussing CO2 elevated levels, it is probably taking a little time to illustrate what makes this area of CO2 occurrence so interesting in terms of the differential levels of reactivity when hydrated versus non-hydrated salt units are intruded and how this process facilitates penetration of volcanic volatiles (including CO2) into such zones. The Herfa-Neurode potash mine is located in the Werra-Fulda Basin in the Hessian district of central Germany (Figure 5a). The targeted ore levels consist of the carnallite-rich Kaliflöz Hessen (K1H) and Kaliflöz Thüringen (K1Th) intervals, which form part of the Zechstein 1 (Z1) bedded Werra salt succession(Warren, 2016). In the mine the K1H and K1Th units range in thickness from 2 m to 10 m, are generally subhorizontal and occur at a depth of 650–710 m below the present-day surface. In the later Tertiary, basaltic melts intruded these Zechstein evaporites as numerous sub-vertical dykes, but only a few dykes attained the Miocene landsurface. Basaltic melt production was related to regional volcanic activity some 10 to 25 Ma. Basalts exposed in the mine walls, where it cuts non-hydrous units of halite or anhydrite, are typically subvertical dykes, rather than subhorizontal sills. The basalts are phonolitic tephrites, limburgites, basanites and olivine nephelinites. Dyke margins are usually vitrified, forming a microlitic limburgite glass along dyke edges in contact with salt (Figure 5b; Knipping, 1989). At the contact on the evaporite side of the glassy rim there is a cm-wide carapace of high-temperature salts (mostly anhydrite and ferroan carbonates). Further out, the effect of the high-temperature envelope is denoted by transitions to clear halite, with higher temperature fluid inclusions (Knipping 1989). All of this metre-scale alteration is an anhydrous alteration halo, the salt did not melt (melting temperature of 804°C), rather than migrating, the fluid driving recrystallisation was largely from entrained brine/gas inclusions. The dolerite/basalt interior of the basaltic dyke is likewise altered and salt soaked, with clear, largely inclusion-free halite typically filling vesicles in the basalt.

    Heating of hydrated salt layers, adjacent to a dyke or sill, tends to drive off the water of crystallisation (chemical or hydration thixotropy) at much lower temperatures than that at which anhydrous salts, such as halite or anhydrite, thermally melt (Figure 5c; Schofield et al., 2014). For example, in the Fulda region, the thermally-driven release of water of crystallisation within particular salt beds creates thixotropic or subsurface “peperite” textures in carnallitite ore layers. These are layers where heated water of crystallisation escaped from the hydrated-salt lattice. Dehydration-driven loss of mechanical strength focuses zones of magma entry into particular subhorizontal horizons in the salt mass, wherever hydrated salt layers were present. In contrast, dyke and sill margins are much sharper and narrower in zones of contact with anhydrous salt intervals and the intrusive is sub-vertical to steeply dipping (Figure 5b versus 5c).

    Accordingly, away from the immediate vicinity of the direct thermal aureole, heated and overpressured dehydration waters can enter a former carnallite halite bed, and drive the creation of extensive soft sediment deformation and peperite textures in the former hydrated layer (Figure 5c). Mineralogically, sylvite and coarse recrystallised halite dominate the salt fraction in the peperite intervals of the Herfa-Neurode mine. Sylvite in these altered zone is a form of dehydrated carnallite, not a primary-textured salt. Across the Fulda region, such altered zones and deformed units can extend along former carnallite layer to tens or even a hundred or more metres from the dyke feeder. Ultimately, the deformed potash bed passes back out into the unaltered bed, which retains abundant inclusion-rich halite and carnallite (Schofield et al., 2014). That is, nearer the basalt dyke, the carnallite is largely transformed into inclusion-poor halite and sylvite, the result of incongruent flushing of warm saline fluids mobilised from the hydrated carnallite crystal lattice as it was heated by dyke emplacement. During Miocene salt alteration/thermal metamorphism in the Fulda region, NaCl-fluids were mixed with fluids and gases originating from thermally-mobilised crystallisation water in the carnallite, as it converted to sylvite. This brine/gas mixture altered the basalts during post-intrusive cooling, an event which numerical models suggest was quite rapid (Knipping, 1989): a dyke of less than 0.5 m thickness probably cooled to temperatures less than 200°C within 14 days of dyke emplacement. The contrast in alteration extent between anhydrous and hydrous salt layers shows alteration effects are minimal wherever the emplacement temperature of the magma is below that of the anhydrous salt body as it is next to a basalt dyke. If this is the mechanism driving entry of igneous-related volatiles (gases and liquids) into a salt body then the distribution of products (including CO2) will be highly inhomogeneous and related to the minerally of the salt unit adjacent to the intrusive.


    Many hydrogen occurrences are co-associated with occurrences of potash minerals, especially the minerals carnallite and sylvite. For example, mine gases (free gas) at Leopoldshall Salt Mine (Zechstein, Permian of Stassfurt, Germany) flowed for at least 4.5 years, producing hydrogen at a rate of 128 cubic feet per day (Rogers 1921). Bohdanowicz (1934) lists hydrogen gas as being present in evaporite intersection in the Chusovskie Gorodki well, drilled in 1928 near the city of Perm to help define the southern extent of the Soligamsk potash. Gases in the carnallitite interval in that well contained 33.6% methane and 17.4% hydrogen. More recent work in the same region clearly shows hydrogen is a commonplace gas in the mined Irenskii unit in the Verkhnekamskoe potash deposit within the central part of the Solikamsk depression in the Ural foredeep. Based on a study of free gas and inclusion-held gas in the Bereznikovshii Mine, Smetannikov (2011) found that the elevated H2 levels are consistently correlated with the carnallite and carnallite-bearing layers (Table 2). Other gases present in significant amounts, along with the hydrogen, in the potash zones include nitrogen and methane. Interestingly, methane is present in much higher proportions in the free gas fraction in the ore zones compared to gases held in inclusions in the potash crystals (Table 2).  

    Smetannikov (2011) goes on to suggest that likely H2 source is via radiogenic evolution of released crystallisation water hence the higher volumes of hydrogen in the carnallitite units in the mine (Table 2). He argues the most probable mechanism generating H2 is the radiolysis of the crystallisation water of carnallite (CaMgCl3.6H2O) driven by the effects of radioactive radiation. The most likely radiogenic candidates are 40K and 87Rb, rather than such heavy radiogenic isotopes as 238U, 235U, 234U, 232Th, and 226Ra. His reasons for this are as follows: 1) U, Th, and partly Ra are sources of α radiation. U, Th, and Ra are concentrated in the insoluble residues of the salts, and the chloride masses contain only minor amounts of Th. Hence these components have no radioactive effect on carnallite because of the short distances of travel of α particles. Because of this, Smetannikov concludes these elements and not likely sources of radioactive radiation. He argues it is more likely that crystallisation water is more intensely affected by β and γ radiation generated by 40K and 87Rb. Hence, bombardment by β and γ radiation drives the radiolysis (splitting) of this water of crystallisation, so driving the release of hydrogen and hydroxyls. Free hydroxyls can then interact with Fe oxides to form hydro-goethite and lepidocrocite, i.e., both these minerals occur in the carnallite but are absent in the sylvinite.

    The notion of hydrogen being created by radiolysis of potash salt layers is not new; it was used as the explanation of the hydrogen association with various potash units by Nesmelova & Travnikova (1973), Vovk (1978) and Knape (1989). Headlee (1962) attributed the occurrence of hydrogen in salt mines to the absence of substances with which hydrogen could react within the salt beds once it was generated. It is likely that there are several different origins for hydrogen gas in evaporites: 1) Production during early biodegradation of organic matter, co-deposited with the halite or potash salts and trapped in inclusions as the crystal grew. This can explain some of the associated nitrogen and oxygen; 2) A significant proportion can be produced by radiolysis associated with potassium salts (when present) and 3) the hydrogen may be exogenic and have migrated into the halite formations, along with nitrogen. 

    Temperature and mineralogical effects on gas generation and distribution in salt (in part after Winterle et al., 2012)

    Temperature can affect brine chemistry of volatiles released as natural rock salt is heated (is this an analogue to the generation of some types of free gas and other volatile released as salt enters the metamorphic realm? –see Warren 2013; Chapter 14). Uerpmann and Jockwer (1982) and Jockwer (1984) showed that, upon heating to 350°C [662°F], the gases H2S, HCl, CO2, and SO2 were released from blocks of natural salt collected from the Asse mine in Germany. Pederson (1984) reported the evolution of HCl, SO2, CO2, and H2S upon heating of Palo Duro and Paradox Basin rock salt to 250°C [482°F]. Impurities within the salt apparently contain one or more thermally unstable, acidic components. These components can volatilize during heating and increase the alkalinity of residual brines. For example, pH of brines increased from near neutral to approximately 10 in solutions prepared by dissolving Permian Basin salt samples that were annealed at progressively higher temperature [up to 167°C [333°F]  (Panno and Soo, 1983).

    Zones of igneous emplacement and intrusion of interlayered halite and potash units create a natural laboratory for the study of the generation and migration of free and inclusion gases during the heating of various salts (Figures 5, 6 and Table 1). In the Cambrian succession of the Siberian platform evaporite intervals are dominated by thick alternating carbonate- sulphate and halite beds. Numerous basaltic dykes and sills intrude these beds. In a benchmark paper dealing with the zone of alteration of intrusives in evaporites, Grishina et al. 1992 found that, in potash-free halite zones intersected by basaltic intrusions, the evolution of the inclusion fluid chemistry is described as a function of the thickness of the intrusion (h) and the distance of the sample from the contact with the intrusion (d) and expressed as a response to the measure d/h. The associated gas in the halite is dominated by CO2 (Table 1). Primary chevron structures with aqueous inclusions progressively disappear as d/h decreases; at d/h < 5 a low-density CO2 vapour phase appears in the brine inclusions; at d/h < 2, a H2S-bearing liquid-CO2 inclusions occur, sometimes associated with carbonaceous material and orthorhombic sulphur, and for d/h < 0.9, CaCl2, CaCl2.KCl and n CaCl2.n MgCl2 solids occur in association with free water and liquid CO2 inclusions, with H2S, SCO, and Sg. The d/h values marking the transitions outlined above occur both above and below sills, but ratios are lower below the sills than above, indicating mainly conductive heating below and upward vertical fluid circulation above the sill. The water content of the inclusions progressively decreases on approaching the sills, whereas their CO2 content and density increase.

    Carnallite, sylvite and calcium chloride salts occur as solid inclusions in the two associations nearest to the sill for d/h<2. Carnallite and sylvite occur as daughter minerals in brine inclusions. The presence of carbon dioxide is interpreted to indicate fluid circulation and dissolution/recrystallization phenomena induced by the basalt intrusions. The origin of carbon dioxide is related to carbonate dissolution during magmatism. Similar conclusions as to the origin of the CO2 in heated halite-dominant units were reached by many authors studying gases in the Zechstein salts in the Werra Fulda region of Germany (Figure 6; Table 1; see Knipping et al., 1989, Hermann and Knipping 1993 for a summary).

    When the gas distributions measured in inclusions in potash units, other than the Cambrian salts of Siberia, are compared to those salts that have not experienced the effects of igneous heating, there is a clear separation in terms of the dominant inclusions gases (Table 1; Grishina et al., 1998). For example, inclusions in the Verhnekamsk deposit (Russian platform) are N2-rich, in regions not influenced by magmatic intrusives (Figure 2, 3). It is an area marked by the presence of ammonium in sylvite (0.01-0.15% in sylvinite and 0.5% in carnallite, Apollonov, 1976). Likewise, nitrogen (via crush release of the samples) is the dominant gas according to the bulk analyses of the same salts by Fiveg (1973). 

    Later Raman studies of individual inclusions in these Cambrian salts reveals a more complicated inclusion story. There are three types of inclusion fill; a) gas, b) oil and c) brine + carnallite-bearing inclusions. Fe-oxides are sometimes associated with inclusions containing the carnallite daughter minerals. Detailed work by Grishina et al. (1998) shows there two kinds of gassy inclusion: 3) N2-rich and 2) CH4-rich 3) CO2-rich in the same age salt (Table 1; Figure 6). That is, not all gassy brine inclusion in the Cambrian salts are nitrogenous. N2 gas inclusions that also contain CO2 and are associated with sylvite with a low ammonium content (0.04 mol% NH4C1). In contrast, CH4 inclusions are associated with ammonium-rich sylvite (0.4 mol% NH4Cl) (Table 2). Older bulk analysis studies(Apollonov, 1976) showed that red sylvinite  has a lower molar NH4Cl content (0.01%) than pink and white sylvinites (0.05 to 0.19%)

    Raman studies of inclusions in the potash-entraining Eocene basin of Navarra, (Spain) outside of any region with magmatic influence show that the gaseous inclusions are mostly N2-rich with 10% to 20% methane (Table 1; Figure 6; Grishina et al., 1998). Traces of CO2 are also detected in some of the Spanish inclusions. Sylvite inclusions in CO2-free inclusions in Spain contain up to 0.3 mol% NH4C1 (Table 2). Grishina et al. (1998) notes that salt formations in the Bresse basin (France) and Ogooue delta (Gabon) have no basalt intrusions and both occur in N2-free, oil-rich environments. The inference is that nitrogen in some salt units is not an atmospheric residual.

    To test if there may be a mineralogical association with a gas composition in inclusions in various salt and evaporitic carbonate layers we shall return to the Zechstein of Germany and the excellent detailed analytical work of Knipping (1989) and Hermann and Knipping (1993). This work is perhaps the most detailed listing in the public realm of gas compositions inclusions sampled down to the scale of salt layers and their mineralogies. Figure 7 is a plot I made based on the analyses listed in Table 9 in Hermann and Knipping (1993).  It clearly shows that for  Zechstein salts collected across the mining districts of central Germany this is an obvious tie of salt mineralogy to the dominant gas composition in the inclusions. In this context, it should be noted that all Zechstein salt mines are located in halokinetic structures with mining activities focused into areas where the targeted potash intervals are relatively flat-lying. There is little preservation of primary chevrons in these sediments. Nitrogen is the dominant, often sole gas in the halite-dominant units, CO2 is dominant in carbonate and anhydrite dominant layers, this is especially obvious in units originally deposited near the base of the Zechstein succession. Hydrogen in small amounts has an association with inclusions the same carbonates and anhydrites, but elevated hydrogen levels are much more typical of potash units, clays and in juxtaposed layers.  

    In my opinion, the gas compositions in inclusions that we see today in any salt mass that has flowed at some time during its diagenetic history will likely have emigrated and been modified to varying degrees within the salt mass. This is true for all the gases in salt, independent of whether the gas is now held in isolated pockets, fractures or fluid inclusions, Non of the gas in halokinetic salt is not in the primary position. Movement and modification of various gas accumulations in halokinetic salt is inherent to the nature of salt flow processes. Salt and its textures in any salt structure have migrated and been mixed and modified, at least at the scale of millimetres to centimetres, driven by vagaries of recrystallisation as a flowing salt mass flows (Urai et al., 2008). All constituent crystal sizes and hence gas distributions across various inclusions in the salts are modified via flow-induced pressure fields, driving pressure solution and reannealing (See Warren 2016 Chapter 6 for detail).

    With this in mind we can conclude that for the Zechstein of central Germany, nitrogen was likely the earliest gas phase as it occurs in all units. On the other hand, CO2, with its prevalence in units near the base of the succession or in potash units that  have once contained hydrated salts at the time of igneous intrusion, entered along permeability pathways. This may also be true of carbonates and anhydrites which would have responded in more brittle fashion. Hydrogen is clearly associated with potash occurrence or clays and an origin via radiolysis is reasonable.

    This leaves methane, which as we saw earlier is variable present in the Zechstein, but not studied in detail by Knipping (1989) or Hermann and Knipping (1993). There is another excellent paper by Potter et al. (2004) that focuses on the nature of methane in the Zechstein 2 in a core taken in the Zielitz mine, Northeastern Germany Bromine values show a salting-upward profile with values exceeding 200 ppm in the region of potash bitterns (Figure 8a). This is a typical depositional association, preserved even though textures show a degree of recrystallisation and implying there have not been massive fluid transfers since the time the salt was first deposited. Methane is present in sufficient volumes to be sampled in the lower 10 metres of the halite (Z2NAa) and in the upper halite (Z2Nac) and the overlying potash (Z2Kst). If was variably present in the intervening middle halite. When carbon and deuterium isotope values from the methane in the lower and upper parts of the stratigraphy are cross plotted. Values from the lower few meters of the halite plot in the thermogenic range and imply a typical methane derived via catagenesis and possible entry into the lowermost portion of a salt seal. The values from the upper halite and the potash interval have very positive carbon values so that the resulting plot field lies outside that  typical of a variety of methane sources (Figure 8b). Potter et al. (2004) propose that these positive values show preserve primary values and that this methane was sealed in salt since the rock was first deposited. That is positive values preserve evidence of the dominant isotopic fractionation process, which was evaporation of the mother brines. This generated a progressive 13C enrichment in the carbon in the residual brines due to preferential loss of 12CO2 to the atmosphere. The resulting CH4 generated in the sediments, as evaporation and precipitation advanced, so recording this 13C enrichment in the carbon reservoir. Therefore, the isotopic profile observed in this sequence today represents a relict primary feature with little evidence for postdepositional migration. This is a very different association to the methane interpretation based on gases held the US Gulf coast or the Siberian salts. 

    The most obvious conclusion across everything we have considered in this article is that, at the level of gas in an individual brine inclusion measure, there is not a single process set that explains gas compositions in salt. Any gas association can only be tied back to its origins if one studies gas compositions in the framework of the geological history of each salt basin. We shall return to this notion in the third article in this series when we will lock at emplacement mechanisms that can be tied to depositional and diagenetic features and compositions at the macro scale.


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    Gases in Evaporites; Part 1 - Rockbursts and gassy outbursts

    John Warren - Monday, October 31, 2016

    The next three articles discuss gases held within salt and is an attempt to address the following questions; 1) What is the scale and location of known rock-bursts/gas-outbursts in salt rock 2) Where do gases reside in a salt mass at the micro- and meso-scale? 3) What are the gases held in salt? 4) How are gassy salts distributed across various salt deposits across the world (macro-scale) and what are the lithological associations? Topics 1 and 2 are the main focus of the first article, topic 3 mostly in the second, while topic 4, where do gases held in salt generate and reside at the scale of a salt mass or salt bed
    is the focus of article 3. Along the way, we shall also discuss whether some of the encapsulated gases in salt can be considered samples of the ambient atmosphere that have been held in brine inclusions since the salt bed was first precipitated? And, as a corollary, we will come to a discussion of how did some of the occluded gases first enter or remobilize through the salt mass during the long history of burial and salt flow (halokinesis) experienced by all ancient evaporite units?

    Gases in evaporites can create problems

    Various gases such as, carbon dioxide, nitrogen, methane, hydrogen and hydrogen sulfide, can occur in significant volumes in and around domal salt masses or bedded evaporite deposits, as seen in numerous documented examples in mines and drilling blowouts in Louisiana, New Mexico, Germany, Poland and China (Figures 1, 2; Table 1). Gases are held in pressurized pockets in the salt that, if intersected, can create stability and safety problems during an expansion of operations in an active salt mine or during petroleum drilling, especially if the pockets contain significant levels of toxic or flammable gases, sufficient to drive rockbursts or gassy outbursts into the adjacent opening. A gas outburst (or rockburst) is defined as an unexpected, nearly instantaneous expulsion of gas and rock salt from a mine production face, normally resulting in an expanded open cavity in the salt. Outburst cavity shapes are generally metre- to tens of metre-scale combinations of conical, cylindrical, hemispherical, or elongated shapes with an elliptical cross section decreasing in diameter away from the opening (Figure 1). Many mapped examples in salt mines of the US Gulf coast have the shape of a cornucopia (Molinda, 1988).

    In the case of blowouts during oil-well drilling, there are two dominant styles of overpressured-salt encounters. The first, and the main focus in blowout discussions this article) is when gassy fluid outbursts occur internally in the salt unit as it is being drilled. Generally, this happens on the way to a test a deeper subsalt target, or less often on the way to test as series of intrasalt beds. Once intersected, pressures in such intrasalt pockets tend to bleed off and so decrease in hours to days as pressure profiles return to normal (Finnie, 2001; Warren 2016; Chapter 8). Providing the drilling system was designed to deal with short-term high-pressure outbursts, drilling can continue toward the target. The other type of gas outburst encountered when drilling salt is located in or near the periphery of a salt mass or bed, especially where the drill bit breaks out on the other side of a salt mass into a highly overpressured and fractured fluid reservoir. Such intersections allow the drill stem to connect with a large highly-overpressured volume of fluids, with the open fractures facilitating extremely high rates of fluid flow into the well bore. This type of breach draws on a significant fluid volume and a resulting blowout can continue unabated for weeks or months.

    Perhaps one the most impressive examples of this type of blowout, and the ability of evaporite unit to seal and maintain an overpressured subsalt pressurized cell, comes from the Alborz 5 discovery in Central Iran (Figure 2; Morley et al., 2013; Gretener, 1982; Mostofi and Gansser, 1957). Earlier wells testing the Alborz Anticline had failed to reach target due to drilling difficulties coming from “an extremely troublesome evaporite section[i] that continually menaced drilling and caused numerous sidetrack operations.” So difficult was drilling through this stressed Upper Red Formation salt unit that it had taken eight months for a previous well to drill through some 170 metres of evaporitic sediments to reach the Qom target. Later wells testing a Qom Fm. target, like Aran-1 to the south of the Alborz anticline, did not intersect thick stressed halite above the Qom Fm., only an anhydrite layer that perhaps was the dissolution residues of former halokinetic salt mass (pers obs.). The discovery well in the Alborz anticline (Alborz 5) had drilled through some 2296 m of middle to late Tertiary clastics and some 381 metres of Oligo-Miocene salines in the lower part of the Upper Red Formation and made up of siliciclastics, banded salt, anhydrite (Figure 3). On its way to the blowout point, the lower part of the well trajectory had penetrated normally to slightly overpressured dirty salt (halokinetic) and then penetrated some 5 cm into the fractured subsalt Qom Limestone (Oligo-Miocene). On August 26, 1956, the entire drill string and mud column were blown back out the hole and many metres into the air. At that time, the mud pressure was 55 MPa (8,000 psi) at a reservoir depth of 2700 m (8,800 ft), a pressure depth ratio of 20.5 kPa/m or 0.91 psi/ft (a lithostatic value!). Over 82 days, the well released 5 million barrels of oil and a large, but unknown quantity of gas before it self-bridged and the flow died on November 18, 1956. The temperature of the oil at the surface was measured at 115°C and at the time of the blowout the mud column density was 2.07 x 103 kg/m3 (129 lb/ft3)(see Figure 3). This type of subsalt overpressured gas occurrence illustrates salt’s ability to act as a highly effective seal holding back huge volumes of highly overpressured fluid. Associated occluding processes are discussed in an earlier series of Salty Matters articles dealing with salt as a seal, especially the article published March 13, 2016.


    Gassy salt (knistersalz)

    Much of the occluded gas in a salt body, prior to release into a mine opening or well bore, is held within inclusions within salt crystals or in intercrystalline positions between the salt crystals. Gas-entraining rock salt, was known from salt mines of Poland and in East Germany since the 1830s and described as knistersalz (literally translates as “crackling salt”). In many mines, walking on knistersalz releases gas as little popping sounds from underfoot. The pressure of the shoe adds a little more stress to an already gas-stressed fragment of salt (Roedder, 1972, 1984). Dumas (1830) first described such “popping salt from Wieliczka, Po­land, and concluded that gas was evolved, presumably from compressed gas inclusions, upon dissolving the salt. Further details on the occurrence were given by Rose (1839). As we shall see, this type of salt can cause serious mine accidents when large volumes of salt explo­sively and spontaneously decrepitate into the mine openings as rockbursts. Dumas (1830) and Rose (1839) found the released gas from "popping " salt in Germany to be inflammable. Bun­sen (1851, p. 251) found 84.6 % CH4 in the gas released during the dissolution of Wieliczka salt, while in many early mines in Germany the occluded gas phase is dominated by nitrogen or carbon dioxide (see Article 2). 

    Knistersalz will "pop" sporadically once placed in water, releasing pressurized gas bubbles as the salt matrix dissolves. This simple demonstration of gas presence is also the foundation for one method of determining the gas content of a rock salt sample (Hyman, 1982). The sometimes rather energetic "pops" that can occur as gases are released from a gas-enriched rock salt sample attest to the high pressures under which the gases are occluded. Pressures postulated in knistersalz can be near-lithostatic and even higher depending on local stresses, related to the low creep limits of rock salt, particularly around mine openings. According to Hoy et al. 1962, CO2-bearing gas mixtures in the knistersalz of the Winnfield salt dome (Louisiana, USA) is under a pressure of 490 - 980 bar (49 - 98 MPa) at 0°C. Similar values (500 - 1000 bar or 50 - 100 MPa) are given by Hyman (1982) for gas bubbles held in rock salt in various Louisiana salt domes. For example, during exploratory drilling in one such Louisiana salt dome, methane gas was released from the salt under a pressure of 62 bar (6.2 MPa) at a flow rate of 1.2 m3/hr (Iannachione et al., 1984). 

    Mining causes a pressure drop in the rock salt as it is extracted from a working face and such pressure drops can change the phase of a fluid occluded in salt, or change the solubility of a gas dissolved in such a fluid. Carbon dioxide, in particular, is susceptible to a phase change because its critical point is close to some ambient mining conditions. As long as CO2 is present above 1070 psi (7.4 MPa) and below 31°C (88°F; critical point), it will be in a liquid phase. Such conditions are not typical in salt mines in the US. However, CO2 generally exists as a liquid in rock salt in many German potash mines (Gimm, Thoma and Eckart, 1966). When mining drops the pressure (from lithostatic to near atmospheric) the CO2 phase will change to a gas, causing abrupt expansion. The sudden change also results in a 5 to 6°C cooling, as measured in regions near large outbursts (Wolf, 1966). The solubility of gases dissolved in brine also changes when mining. For example, the solubility of methane in brine is extremely low at atmospheric pressure and so is released as gas bubbles from a brine issuing from rock salt fissures upon mining, as observed in a number of US Gulf Coast salt mines (Iannacchione and Schatzel, 1985).

    Pressures released during an outburst result in velocities at the outburst throat which can be very large and locally approach sonic velocities (Ehgartner et al., 1998). Velocities of more than 152 m/sec (500 ft/s) have been recorded in vertical airways some distance from rockbursts in Germany. Velocities at the rockburst site would be even higher. Narrow throat characteristic of some rockbursts can result in throttling. However, associated pressure waves are not strong enough to cause the observed levels of equipment destruction, since they are of a magnitude similar to those found in blasting. Rather, observed damage associated with rockbursts is due to flying debris in the pressure wave as the quantities of rock thrown out by the burst have high kinetic energy (Wolf, 1966). 

    Given the relatively impermeable nature of bedded and halokinetic salt, occluded gases generally are not released from their containment unless mining or drilling activities intercept (1) a gas-filled fissure zone, an area where the voids between the salt crystals are interconnected, (2) a mechanically unstable zone of gas-enriched salt that disaggregates, releasing its entrained gases (a blowout), or (3) as the mine or the drill bit enters some other relatively permeable geologic anomaly (Kupfer, 1990).


    Gassy outbursts and rockbursts in salt

    Outbursts are documented in the U.S., Canada, and throughout northern Europe in various salt and potash mines (Figure 2; Table 1). The salt domes of northern Europe and the US Gulf coast are in particular loaded with pockets of abundant gas inclusions (Ehgartner et al., 1998). Many dangerous pockets of methane and H2S were intersected during the opening of shafts into the domes of Zechstein salts in the Saxony region, Germany and several early potash mines in the area were abandoned because of problems caused by rockbursts and associated gas outflows (Gropp, 1919; Löffler, 1962; Gimm, 1968). Before the current practice of evacuating any gas-prone salt mine prior to blasting, many fatalities resulted from such gas and rock outbursts (Table 1). A significant portion of the deaths was due to secondary factors (post-rockburst), such as methane fires, CO2 suffocation, and H2S poisoning (Dorfelt, 1966). Even with the practice of mine evacuation prior to blasting, outburst gases have in some cases filled a mine, blown out of the mine shafts, and caused fatalities at the surface. This was the case in Menzengraben in 1953, as heavier-than-air CO2 gas, released by a blasting-induced rockburst, blew out of the mine shafts for 25 minutes and flowed downhill into a nearby village, where it ponded and ultimately suffocated 3 people in their sleep (Hedlund, 2012)

    The most frequent and largest rockbursts and gas outflows from subsurface salt occurred in the Werra mining district in former East Germany. Gimm and Pforr (1964) report that rockbursts occurred every day in the Werra region. If one also includes potash mines in the Southern Harz region, more than 10,000 outbursts were recorded up till the 1960s in the German salt mines (Dorfelt, 1966). The 1953 Menzengraben(Potash Mine No. 3) rockburst blew out some 100,000 metric tons of fractured rock salt (approximately 1.6 million cubic feet). This may well be the world’s largest rockburst in terms of cavity size (Gimm, 1968). In an earlier incident in the same region in 1886, the shaft Aschersleben II was flooded with water and gas as it reached a depth of 300 m. A pilot hole drilled from the temporary bottom of the shaft into the underlying Stassfurt rock salt, hit a gas pocket, releasing a combination of H2S—CH4—N2 gases, which then escaped under high pressure for some two hours carrying with it an NaCl brine to the height of a “house” above the shaft floor before the outflow abated. The shaft was abandoned (Baar, 1977).

    In 1887 the shaft Leopoldshall III, at Stassfurt, had been sunk through the caprock, and into the Zechstein salt to a total depth of 412 m subsurface, when it hit a gas pocket containing H2S, and four miners were killed by gas escape. Subsequently, in 1889, seven more were killed during shaft construction in the same mine. In 1895, a large volume of CO2 was released from rock salt at a depth of 206 m during the sinking of the Salzungen shaft (Gimm 1968, p. 547). Numerous other outbursts of gas occurred in the same Werra-Fulda district with most mines operating at depths greater than 300 meters, with outbursts responsible for a number of deaths both below and above ground. According to Gimm (1968, p. 547), since 1856, toxic gases were also encountered during the sinking of a number of other shafts in the Stassfurt area. Gropp (1918) documents 106 gas occurrences in German potash mines for the period 1907 to 1917, at depths of ≈300 meters and greater. Many of these gassy encounters caused casualties, particularly in salt dome mines of the Hannover area where several of the potash mines were abandoned due to dangerous gas intersections (Barr, 1977).

    Less severe examples of gas outbursts and rockbursts transpired in other salt mines around the world (Figure 2). More than 200 gas outbursts with ejected rock salt volumes up to 4500 tons have occurred in the Upper Kama potash deposits of Russia (Laptev and Potekhin, 1989). Baltaretu and Gaube (1966) reported sudden gassy outbursts in potassium salt deposits in Rumania. Outbursts in Polish salt mines were noted by Bakowski (1966). Potash mines in England and Canada also exhibited outbursts (Table 1; Schatzel and Dunsbier, 1988) with the most recent case being a gassy outburst that caused a fatality in the Boulby mine in July 2016.

    Major rockbursts, tied to methane releases, occurred in Louisiana in four of the 5-Island salt mines exploiting the crestal portions of subcropping salt domes (Belle Isle, Cote Blanche, Weeks Island, and Jefferson Island) with the exception of Avery Island. Gassy outbursts, of mostly CO2, also occurred at the Winnfield salt mine, Louisiana (Table 1). Rockburst diameters range from a few inches up to over 50 ft. Cavity heights range from several inches to several hundred feet. Smaller rockburst and cavities in the Five-Island mines were ordinarily not reported (Kupfer,1990). Only the more gas-inclusion-rich salt decrepitates in these mines, and the concave curvatures of the walls are such that the resulting slight additional confining force from the concavity keeps the remaining salt from decrepitating further (Figures 1, 4; Roedder, 1984).

    The larger outburst shapes tended to be cornucopian in shape, whereas the shorter ones were conchoidally shaped with symmetrical dimensions (Figure 4). Outbursts approaching several hundred feet high were documented in the Jefferson Island and Belle Isle mines. The disaster at Belle Isle mine in 1979, in which five miners died, proved that high-pressure methane in large quantities could be released near instantaneously during a rockburst. It was estimated that more than 17,000 m3 (600,000 ft3) of methane was emitted by the 1979 outburst (Plimpton, et al.,1980). At the former Morton mine at Weeks Island, an even larger gas emission apparently occurred in connection with a rockburst. It was estimated that as much as 1,020 m3 (36,100 ft3) of salt was released as 1.4 million m3 (50 million ft3) of gas filled the former Morton Mine (MSHA,1983). If the limited number of sample points represent a well-mixed mine atmosphere, the gas alone would occupy approximately 17,000 m3 (600,000 ft3) in the salt at lithostatic pressure (Plimpton, et al.,1980).

    Outbursts occurred during mining in all three of the mines at Weeks Island - the “old” Morton mine (the site of the now abandoned U.S. Strategic Petroleum Reserve), the Markel mine, and the “new” Morton mine. Perhaps the largest outburst at the “new” Morton mine occurred on October 6, 1982, in the southwest corner of the 1200-ft level, close to the edge of the dome. A balloon with an attached measuring string is typically used to estimate the height of the major vertical outbursts. A balloon went up more than 30 m (100 ft) into an outburst some 10 m (35 ft) wide (MSHA, 1983). Outbursts in the old Morton mine occurred only in the larger lower level (-800 ft) of the two level mine outside the vertically projected boundary of the upper (-600 ft) level. A similar trend was noted at Jefferson Island where no gas outbursts occurred in the upper level of the mine. The outbursts observed at the Jefferson Island mine were in the same relative position at both the 1300-ft and 1500-ft levels. This is attributed to the near vertical orientation of a very gassy zone of salt (Iannacchione, et al., 1984). Structural continuity (banding) is nearly vertical in many Gulf coast salt dome diapirs, except where the top of the dome has mushroomed. As a result, horizontal runs of outbursts have reportedly been small, and unlikely to connect caverns separated by 100 ft or more (Thoms and Martinez, 1978.).

    The geometry of the gas pockets is not well known. Thoms & Martinez (1978) argued that prior to the rockburst the gas is concentrated in vertical, cylindrical zones or pockets, which were created and elongated by the upward movement of the salt. Mapping in the Five-Island mines shows that the rockbursts are often aligned along structural trends . At Winnfield (Hoy et al., 1962), and possibly at Belle Isle (Kupfer,1978), the outbursts occur close to the edge of the dome. In other cases (e.g., Cote Blanche and Belle Isle) the outbursts follow structural trends such as shear zones within the dome (Kupfer, 1978). In all cases, there is an association between methane gas occurrence and other anomalous features such as dirty salt, sediment inclusions and oil or brine seeps (see article 2).

    Rockbursts are not limited to gassy intersections in domal salt. High-pressure pockets of inert gas, typically nitrogen, are documented in bedded potash mines (Carlsbad, NM), and combustible gases (methane)and fluids (brine and oil) in potash mines in Utah (Djahanguiri, 1984). The Cane Creek potash mine (Utah). exploiting halokinetic salts sandwiched by the bedded formations of the Paradox Basin, had a history of fatalities and extensive equipment damage as a result of rockbursts (Westfield, et al., 1963). In contrast, no gassy outbursts were reported during the construction and operation of the Waste Isolation Pilot Plant in the bedded salts of southeastern New Mexico. During WIPP construction, routine drilling ahead of the road-header checked for gas, but found very little (Munson, 1997).

    In my opinion, some gas pockets in domal salt can be related to the diagenetic process creating a caprock, where metahne and H2S are typical byproducts. In others, the gases are related to the burial history and recrystallisation (partially preserving primary nitrogen), while in yet others, the gas release is related to heating and alteration especially of the hydrated salts (hydrogen) and associated fracturing related to igneous intrusion (CO2). In some cases, gases were encountered in fracture systems of cap anhydrite close to the top or edge of the salt dome; such fracture systems apparently had connections to the groundwater as the gassy outbursts were followed by water of varying salinity. In other cases, fracture systems headed by a gas cap connected the expanding mine to overlying aquifers and ongoing salt dissolution was facilitated. But, in most cases of rockburst located within the interior of a salt mass, the majority of the intersected gas pockets are isolated, as once the burst occurred most cavities tended to receive little if any subsequent recharge, so gas and brine outflow rates tended to decrease to zero across hours to days (Loffler, 1962). The relationship between the type of gas, its position in the salt, and possible lithological associations are documented and discussed in detail in articles 2 and 3.


    The physics that drives rock and gas outbursts in an expanding mine-face or shaft is relatively straightforward. In the petroleum industry, it constitutes a process set that is already well documented as the cause of many salt-associated gassy blowouts such as Alborz 5 (Figure 3; Warren, 2016 – Chapter 8 for detail on pressure distribution in and about a salt mass). Oilfield blowouts associated with salt occur when pore pressures in fluids in the drilled rock approach or even exceed lithostatic and the weight of mud in the approaching borehole is not sufficient to hold back this overpressured fluids entering and escaping up the borehole (Figure 3). Spindletop and other famous caprock blowouts in the early days of salt dome drilling in Texas and Louisiana are famous examples of this process (Figure 5). Ehgartner et al. (1998) argue that the same pressure release occurs as an expanding mine face approaches a gassy zone in the mined salt. Once the pressure is reduced by the approach of the mine face, the release of gas formerly held in place by lithostatic pressure within a homogenously stressed salt mass will release, the enclosing rock salt will lose cohesion and so a rockburst (gas outburst) occurs (Figure 6).


    How is the gas held and distributed within salt at the micro and mesoscale (microns to metres)?

    That free gas and gas in inclusions occur simultaneously in salt masses is undeniable, numerous examples come from salt mines and salt cores (Table 1). Gases are held in evaporite salts in three ways (Hermann and Knipping, 1993); 1) Crack- and fissure-bound gases, 2) Mineral-bound gases, a) intracrystal, b) intercrystal, and 3) Absorption-bound gases. Type 1 occurrences, as the name suggests, are defined by gas accumulations in open fractures and fissures, typically in association with brine. Some occurrences are tied to pressurized aquifers, others are isolated local accumulations within the salt. Intracrystal gas occurs as bubbles, some elongate, some rounded in brine inclusions that are fully enclosed within a crystal (typically halite). At the micro (thin section-SEM scale), intracrystalline gases typically form as a few to aggregates of small bubbles, arranged along crystallographic axes or planes, with bubble diameters in the range 1 to 100 µm. Intercrystalline gases occupy the boundary planes of crystals in contact with one another, that is intercrystalline gases occupy polyhedral porosity. According to Hermann and Knipping (1993), up to 90% of the mineral-bound CO2gas mixtures in the salt rocks of the Werra-Fulda mining district is likely intercrystalline, and the remaining 10% is intracrystalline. Absorption bonding is likely an independent form of gas fixation in salt. Adsorptive bonding describes the ability of solids, especially clays, and crystalline compounds to store gas on their surfaces in the form of layered molecules, most would term this a subset of microporous gas storage in a shale.

    [i]The stresses in and around and in salt structures can be high and troublesome to stabilize, even today and is an indication of the ongoing dynamic nature of salt flow and recrystallisation in the subsurface.Therefore, if borehole fluid pressure is lower than salt strength during drilling, stress relaxation may significantly reduce open-hole diameters. In some cases, relaxation causes borehole restrictions even before drilling and completion operations are finished and casing has been set.


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    Silica recycling and replaced evaporites - 4. Proterozoic atmospheric transitions and saline microporous chert reservoirs

    John Warren - Friday, October 07, 2016

    Sulphate and oxygenation levels across the Proterozoic

    Proterozoic (2.5 – 0.542 Ga) saline sediments encompass significant transitions in evaporite style and chemistry within an evolving atmospheric and oceanic framework. Lithospheric changes tie to a cooling and biologically-evolving earth as earth-scale plate tectonics move to a system set comparable with that operating today. The Proterozoic eon is divided into three eras: the Paleoproterozoic, Mesoproterozoic and Neoproterozoic. Thick sequences of halite are only found as the actual bedded salts in sediments of the Neoproterozoic (and the Phanerozoic), while calcium sulphate residues and beds occur in all three Proterozoic eras, especially in parts of the Paleoproterozoic and the Neoproterozoic.

    Paleoproterozoic era sediments (2.5-1.6 Ga) contain isotopic evidence the first significant oxygenation event in the world's atmosphere, largely driven by the increasing dominance of cyanobacterial photosynthesis. Neoproterozoic sediments (1.0 – 0.542Ga) contain evidence of the second oxygenation event, which is associated with the evolution of widespread multicellular life and CaCO3/siliceous carapaces. By the end of the Neoproterozoic, the world oceans had chemistries, temperatures and salinities similar to those of the Phanerozoic (Blamey et al., 2016). The intervening Mesoproterozoic (1.6- 1.0 Ga) retains evaporitic residues with aspects of both the late Archean and the Phanerozoic.

    The oxygenation of Earth’s atmosphere-ocean system occurred in two steps: 1) the Paleoproterozoic “Great Oxygenation Event” (GOE ≈ 2.3 Ga), which refers to the transition from a pervasively reducing Earth-surface system to one with an oxygenated atmosphere and oxygenated shallow seas, and 2) the “Neoproterozoic Oxygenation Event” (NOE), when the Earth’s atmosphere and ocean are understood to have become persistently oxygenated down to the deep ocean bottom (Turner and Bekker, 2016; Scott et al., 2014). The GOE is indicated by a Proterozoic carbon isotope anomaly known as the “Lomagundi event,” a positive carbon isotope excursion between ca. 2.22 and 2.06 Ga, interpreted to be the result of high organic carbon burial and attendant accumulation of atmospheric oxygen (Figure 1; Bekker and Holland, 2012)). A long interval spanning the remainder of the Paleoproterozoic and much of the Mesoproterozoic followed the Lomagundi event, typified by lower levels of atmospheric oxygen and little variation in carbon isotope values. This ended in the late Neoproterozoic with dramatic fluctuations, of escalating magnitude, in the biogeochemical carbon cycle and attendant fluctuations within an overall increasing oxygen content (Figure 2). By the end of the Neoproterozoic, not just shallow shelf waters but much of the deep-ocean water column was consistently oxygenated (see article 1 in this series)


    Chert evolution

     We have already seen how low levels of oxygen and high levels of CO2 in the Archean favoured the precipitation of nahcolite, and its hydrothermal silica association, atop subsealevel isolated saline sumps in microcontinents and island arcs in a saline waterworld (article, 28 August 2016). The hydrothermally-dominated silica of the silica-rich Archean oceans is reflected in more negative 30Si isotope values in widespread marine cherts of that time, compared with most Proterozoic cherts (Chakrabarti et al., 2012; see also Figure 3b). However, shorter-term fluctuating levels of atmospheric oxygen in the Proterozoic also influenced drop-out salinities for gypsum in Mesoproterozoic marine brines. In brines derived from the modern, well-oxygenated world oceans, as in evaporite successions deposited throughout the Phanerozoic, the sulphate minerals gypsum and anhydrite precipitated from evaporating seawater after aragonite or calcite, but before halite (see 26 August, 2015, blog for more detail). At lower seawater sulphate levels across the Archean and much of the Proterozoic, gypsum and anhydrite precipitated after halite, even at Na and Cl concentrations similar to those of the modern ocean.

    This is why some post-Lomagundi, Paleoproterozoic marine evaporite successions show clear evidence of halite precipitation before gypsum or anhydrite or even an absence of gypsum or anhydrite with halite (e.g., ≈ 1.88 Ga Stark Formation; Pope and Grotzinger, 2003). The post-halite precipitation of calcium sulphate is construed as evidence for a limited marine sulphate reservoir and little atmospheric oxygen (Scott et al., 2014). In contrast, Lomagundi- age sedimentary successions contain evidence for sulphate precipitation before halite (Melezhik et al., 2005; Bekker et al., 2006; Schröder et al., 2008). Sulphur isotope values of marine sulphates (in CaSO4, barite) and sulphides in marine pyrite also record expansion and contraction of anoxic oceanic settings. That is, a higher burial rate of pyrite in anoxic settings is indicated by a positive shift in the sulphur isotope values of sulphates, whereas ocean oxygenation creates a negative shift in values (e.g., Claypool et al., 1980; Strauss, 1997). Furthermore, expansion of the area of anoxic oceanic settings decreases the size of the seawater sulphate reservoir, resulting in more variable sulphur isotope values of sulphate evaporites, barites, and other carbonate-associated sulphates (Figure 2; Kah et al., 2004). This applies in particular in the Mesoproterozoic when only tshallow oceanic waters were consistently oxygenated.

    When we look at silica mobility and chert styles across a Proterozoic milieu of evolving oxygen and sulphate levels we see some aspects similar to the Phanerozoic and others more akin to the high-silica oceans of the Archean. Maliva et al. (2005) and Perry and 2014 show that the latter part of the Paleoproterozoic era (post-Lomagundi) is marked by the end of widespread primary and early diagenetic silica precipitation in normal marine subtidal environments. However, silica precipitation continued apace in the deeper marine in waters that were still anoxic. The Paleoproterozoic is defined by the “rusting” of the shallower parts (shelves and upper slopes) of the world’s ocea,n as dissolved oxygen levels increased and the accumulation of widespread Banded Iron Formations (BIFs) occurred, including the huge deposits of NW Australia.

    So where and when do we see nodular cauliflower chert after sulphate in the Proterozoic?

    Some of the oldest silicified nodular sulphates with cauliflower chert textures and actual relict anhydrite are found in the Huronian Gordon Lake Formation (≈2.4 Ga; Chandler, 1988). The nodules are commonest in mud chip breccia at the base of sandstone, siltstone and mudstone upward-fining storm cycles. Anhydrite nodule relicts are composed of a mosaic and meshworks of blocky crystal laths. Earlier laths are preserved in silicified outer rims of many cauliflower chert nodules, with texture alignments similar to that of Recent displacive sabkha anhydrite nodules. Completely silicified nodules are composed of megaquartz, some calcite-cored, or of jasper, with replacement textures identical those documented the Phanerozoic by Milliken (1979). Similarly-textured cauliflower cherts are found in the Mallapunyah Formation (1650 Ma) in the Paleoproertozoic sediments of the McArthur Basin in the Northern Territory of Australia (Warren, 1999). Thin sections through 5-30cm diameter Mallapunyah nodules (in a redbed host) can still retain small relict highly-birefringent laths of anhydrite, but most of the former felted cores to the nodules are now composed of mimetic silica. There are also older sedimentary chert nodules in the McArthur Basin succession located in units below the level of the Mallapunyah, but they are smooth walled not rugose-surfaced features.

    Unfortunately, the term cauliflower chert is loosely defined and is used to describe cemented features in Meosproterozoic and earlier sediments and metasediments, which are not true calcium sulphate evaporite replacements.Although termed cauliflower features, they do not have surface textures resembling the florets of a cauliflower (see article 2 in this series - July 31, 2016). For example, aggregates and clusters of growth-aligned barite crystals in the Archean of South Africa are described as cauliflowers when they should be described as bladed, palmate crystal aggregates (Reimers and Heinrich, 1997). Interstingly, Chowns and Elkins (1974) in a study of cauliflower cherts occurrences across the USA list no examples older than Cambrian. Using a tighter definition of cauliflower chert and recognising that this term should not be interchangeable with crocodile-skin chert it seems that Proterozoic occurrences of cauliflower chert nodules largely mirror times when oxygen levels were sufficiently high in the world's ocean to allow sulphate in solution. In the Paleoproterozoic and Mesoproterozoic only the upper parts of the ocean column, including waters covering the world's continental shelves (and derived evaporite basins) were sufficiently oxygenated to allow the formation of cauliflower chert after nodular anhydrite. However in some Neoproterozoic basins, especially if located in sumps in a highly-restricted brine layered seafloor, the levels of anoxia in the ponded bottom brines facilitated the accumulation of laminar microporous chert in association with evaporites or their early replacements

    Primary laminated hypersaline silica chert in an evaporite basin at the Precambrian-Cambrian boundary

    An organic-rich laminated porous chert known as the Athel or Al Shamou silicilyte consists of up to 90% microcrystalline quartz along with dolomite, magnesite, anhydrite and halite (Rajaibi et al., 2015). It occurs at the Precambrian-Cambrian boundary in the subsurface of the South Oman Salt Basin, Sultanate of Oman, where it acts as a light-oil reservoir  (Ramseyer et al., 2013; Amthor et al., 2005). Fully encased in variably halokinetic salt masses, it was first discovered during the 1990's hydrocarbon exploration activities of Petroleum Development Oman. This laminated microporous and variably fractured chert, has its source of silica and its mode of precipitation tied to an anoxic, sulphur-rich, stagnant and highly saline basin. Its homogeneous silica distribution and high Si isotope values (avg. d30Si = +0.83 ± 0.28), coupled with a low molar Ge/Si ratio (<0.25 x 10-6) in its microcrystalline quartz matrix imply dissolved silica in concentrated seawater as the Si source, and hydrothermal or biogenic (e.g. sponge-derived) silica are excluded.

    Silica precipitation from a seawater-sourced brine was likely the result of a dramatic increase in salinity in response to halokinetic salt dissolution atop and adjacent to the edges of transtensional depressions on a deep basin floor in the South Oman Salt basin, thus markedly reducing the solubility of amorphous silica in these brine-filled seawater depressions. This saturation triggered the formation of silica gel. The gel accumulated at the base of a brine-layer covered basin floor, forming a soft silica-rich layer bound into bacterial mats, giving rise to its fine-scale lamination. The mean number of laminae in this laminated chert is ca. 32 per year suggesting that layering is non-annual and controlled by processes such as fluctuations in nutrient supply, lunar driven re-mixing or diagenetic segregation. The transformation of the silica-gel to microcrystalline quartz occurred below 45°C indicating a less than -4.5‰ d18O composition of the pore-water during microcrystalline quartz formation. The  microporous hydrocarbon filled nature of this ancient chert and the fact the hydrocarbon-filled micropores are still distinct after more than 500 million years after they filled (Figure 4d, e) is  why when artificially fractured the silicilyte can act as a hydrocarbon reservoir (See Rajaibi et al. 2015 and Warren, 2016; Chapter 10 for a summary of relevant literature).


    What does this mean for other evaporite-associated laminar cherts in the Phanerozoic?
    Beds of laminated chert (not replaced, but precipitated, primary silica accumulations) are unusual across the Phanerozoic, but cauliflower cherts as replacement of calcium sulphate evaporites are not (Chowns and Elkins, 1974). Spanning the Precambrian boundary, the Athel Silicilyte is the first Phanerozoic example of a laminar chert-hypersaline association. Then, there is the somewhat younger but world-famous Devonian to early Carboniferous Caballos Novaculite. This chert location was the site of a world-famous set of arguments as its origin, between two world-renown professors, Dr Earle McBribe and Dr Robert Folk, both on the faculty of the University of Texas at that time. Folk argued for a shallow-water peritidal hypersaline depositional setting, McBride for a deep a marine setting but still with possible hypersaline indicator textures (Folk, 1973; Folk and Mcbride, 1976; McBride and Folk 1977).

    The Caballos Novaculite outcrops in the Marathon Uplift of Texas, while its lithologic and time equivalent, the Arkansas Novaculite, outcrops in the Ouachita Mountains of Arkansas and Oklahoma (Figure 5). Novaculite chert) in outcrop is very resistant to erosion so that layers of novaculite stand out as characteristic ridges and dip slopes in the Ouachita and Marathon mountains (Figure 5). This outcrop forms and its hard abrasive nature gives it the name novaculite, which in its Latin root novacula, means razor-stone. When some novaculite is fractured in the subsurface, there is sufficient connected porosity to form a fractured reservoir play, as in Arkansas and Texas. There, some 30 years ago, oil and gas fields such as Isom Springs in Oklahoma and McKay Creek, Pinion and Thistle fields in West Texas were discovered in the Caballos and Arkansas novaculite-chert. The chert reservoir is most productive when it is highly fractured, occurs within complex thrust faults and has had any enclosed carbonate material leached from its chert matrix, so creating microporosity (Figure 6; Godo et al., 2011).

    Chert beds in the Caballos Novaculite are composed of equant grains of microcrystalline quartz, minor amounts of illite and radiolaria, and trace amounts of pyrite, carbonate and other minerals and organic matter. The chert beds are generally interpreted as having formed by the silicification and alteration of precursor sediment, sometimes massive, other times finely laminated. Some beds retain occasional evidence biogenic silica derived from radiolaria, while underlying levels retain can contain abundant siliceous sponge spicules. Fractures and crackle breccias developed in a grey chert following lithification; green siliceous sediment, whose lithification was impeded by clay, filled these pre-orogenic fractures.
    Beds of red shale, chert pebble and cobble conglomerate, sandstone, limestone, dolomite, and lumpy manganiferous and jasperitic chert make up no more than 3% of the chert and shale members of the Caballos, but still are of controversial origin and environmental significance. The chert conglomerate beds, for example, are interpreted as tidal-channel deposits by Folk and as mass-flow deposits by McBride. Jasper beds texture are considered bizarre by many geologists who have worked on them: they are lumpy, uneven beds 0.2 to 2 m thick composed of cherry-red chert with local geopetal cavities, contorted laminae. manganiferous zones, cauliflowerlike quartz-filled nodular cavities sometimes with hollow centres and variably filled or partially filled with zebraic chalcedony, lutecite, quartzine, pseudocubic quartz crystals, and filamentous structures resembling algae. These beds are interpreted by Folk as the product of diagenetic alteration of sabkha evaporite nodules and siliceous ooze, during and following subaerial exposure with soil development, and by McBride as the product of diagenetic alteration of evaporite beds deposited in deep water and sandwiched between radiolarian ooze. Synthesis of evidence on the origin of both the novaculite and chert and shale members leads to contrasting interpretations of water depth during deposition. However, if the evaporite solution breccias, recognised as such by both authors and the cauliflower cherts as replaced diagenetic anhydrite clusters then the depositional setting is akin to that of the Athel Silicilyte (namely a deepwater holokinetic hypersaline evaporite setting not unlike a siliceous DHAL association (see Warren, 2016; Chapter 9, for discussion of the DHAL literature).
    Economic implications of understanding what defines a silicilyte versus novaculite versus tripolite versus diatomaceous ooze
    We have already seen that microporous cherts when fractured are possible reservoir rocks as illustrated by the saline-associated Athel Silicilyte and the Caballos Novaculite. The former retains its microporosity because of an early hydrocarbon charge into existing microporosity (Rajaibi et al. 2013; Amthor et al., 2005), while the reservoir quality of the latter was enhanced by diagenetic leaching of finely dispersed carbonate material, likely when it was caught up in the Ouachita Orogeny (Figure 6; Godo et al., 2011). Like the Athel, the Arkansas (Caballos) Novaculite is thought to be self-sourcing in terms of reservoir hydrocarbons (Zemmels et al., 1985).
    The term silicilyte is defined by Rajaibi et al. (2013) as a "locally-used" term to describe porous organic-rich laminated chert, it is a succession of microcrystalline quartz that is preserved within salt-encased slabs, 300 to 400 m thick, at a depth of 4 to 5 km in the South Oman Salt Basin. As mentioned earlier, the term novaculite comes from its outcrop expression and its "razorstone" properties and does not necessarily have a direct connection to microporosity in the reservoir portions of the unit (Figure 5). Outcrops of weathered microporous chert zones in the upper part of the Arkansas (Caballos) Novaculite are called tripoli or rottenstone (Figure 6). When present in this finely powdered microporous form, it is quarried and crushed for use as a polishing abrasive in metalsmithing and woodworking. When present as a very hard dense rock, it can be cut and shaped for use as a whetstone or razorstone. Before European settlement, novaculite was a source for numerous arrow-tips, spear tips and knives. 
    Etymological variations of the terms tripoli, novaculite and silicilyte as forms of chert, as currently used in the geological literature are interesting and geologically confused. This is particularly true if the writer did not 1) understand there are various origins to laminar sometimes microporous cherts, and  2) that there various possible silica sources and precipitation/replacement mechanisms can be halotolerant bacteria, or other silica sources that can be tied to marine sponges and yet others to radiolaria and diatomaceous oozes. Hence, there is the time-related aspect of biogenic chert evolution tying back to the source of the silica in some cherts and the presence or lack of salinity indicators in a laminar chert and chert nodules (e.g., cauliflower versus crocodile-skin cherts).
    A diatomaceous ooze is a form of opaline silica made up of accumulations of siliceous frustules of diatoms in normal marine pelagic sediments and when it retains microporosity is sometimes called tripolite or tripolitic earth. Diatomaceous oozes, the precursor to this type of tripolitic earth is often laminated, but is harsh to the feel and scratches glass. To add to the confusion there are microporous mesohaline diatomaceous oozes called the Tripoli unit in many Messinian subbasins.
    Tripoli or tripolite powder is also the term used to describe the form of microporous laminar chert used as an abrasive and collected from weathered zones of the Caballos and Arkansas Novaculites. The Palaeozoic age of the Caballos and Arkansas Novaculite means it cannot contain diatoms. Diatoms evolved in the Cretaceous and have been the dominant source of remobilised silica in marine chert nodules ever since. Palaeozoic silica remobilized into chert is often related to nearby sponge spicule horizons, or less often to radiolarian beds. The further back in time, and perhaps the more saline the marine bottom water, the greater the likelihood of a microbial association with chert precipitation and remobilisation.
    Even in Tertiary strata some microporous diatomaceous earths can have a normal-marine, organic-enriched depositional association and can constitute a fractured microporous hydrocarbon reservoir. This is the case with the Miocene-age fractured microporous chert reservoirs that produce today in the Santa Barbara Basin of offshore California (Reid and McIntyre, 2001). Cores and coastal outcrops of the Monterey Formation show this type of marine-deepwater diatomaceous ooze is interlayered with microbial (bacterial-archeal) methanogenic organic-rich dolomites. Then, there are the deepwater diatomaceous saline oozes in the Miocene units that immediately underly Messinian evaporites in the Lorca and other sub-basins across the Mediterranean (Rouchy et al., 1998). These diatomaceous organic-rich oozes were deposited as pelagic sediments in saline waters on stratified bottoms that herald the creation of saline bottom water layers related to the onset of hypersaline conditions. Their depositional setting is in restricted basins with increasingly saline bottoms, driven by tectonic isolation and drawdown that soon after precipitated the salts of the Messinian Salinity Crisis. Finally, there are the lacustrine diatomaceous oozes accumulating at the base of density-stratified water columns in lakes of African Rift Valley. This type of laminar ooze occurred in the deeper parts of the lake floor and in Lake Magadi and Lake Natron define units that immediately predate significant lake drawdown episodes that are defined by the type 1 (sodium bicarbonate) evaporite layers. These laminar diatomaceous chert beds contain nodules with the characteristic surface shrinkage textures of crocodile-skin chert (see article 1 in this series of four).
    So, in terms of silica with an evaporite association, "one size does not fit all.' There are multiple saline settings, both depositional and diagenetic, with silica sources evolving with life across the Proterozoic into the Phanerozoic. Interactions between biology and brine chemistry control the accumulation of silica in various forms in a range of evaporite settings, ranging across the marine to deep halokinetic seafloors to lacustrine basins. Once we understand how to recognise cauliflower chert and its possible association with laminar saline cherts of the Proterozoic and the Phanerozoic, then particular chert styles can help to define the evolution of atmospheric oxygen and the saline versus non-saline origins of some organic-rich laminar biogenic microporous cherts.



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    Silica mobility and replaced evaporites: 3 - Archean cherts

    John Warren - Sunday, August 28, 2016


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

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

    Silica levels in the Archean ocean

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


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

    Archean crustal tectonics and silicification of world-scale evaporites

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

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

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

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

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

    Archean oceans and silicified sodic evaporites 

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

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


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

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

    Archean evaporite deposition and silicification

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

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

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

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

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

    Is the present the key to the past?

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

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

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

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



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    Silica mobility and replaced evaporites: 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.


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    Silica mobility and replaced evaporites: 1 - Alkaline lakes

    John Warren - Saturday, July 02, 2016


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


    Silica geochemistry

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

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

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

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

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

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


    Modern siliceous sediments

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

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


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

    Dissolved silica and biogenic silica cycling in mesohaline lakes

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

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

    Biogenic lacustrine silica

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

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

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

    Silica in alkaline surface waters

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

    Coorong cherts

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

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

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

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

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

    Lake Magadi chert

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

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


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

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

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

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


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

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

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

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

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

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

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

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

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

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

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

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


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


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