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 (

Non-solar thick salt masses: Part 2: Oceanic ridge anhydrite and mantle-derived halite

John Warren - Sunday, June 16, 2019


The previous article in this discussion of significant salt volumes not created by solar driven evaporation focused on a number of processes that drive crystallisation, namely temperature changes via brine warming (prograde salts) or cooling, especially cryogenesis, as well as brine mixing. In this article, we shall further develop the notion of temperature changes driving salt crystallisation, but now focus into higher-temperature subsurface realms generally flushed by igneous and mantle fluids.

Most of the precipitates can be considered hydrothermal salts, which is a broader descriptor than burial salts (Warren 2016; Chapter 8), that encompasses a higher temperature range compared to the diagenetic realm. One group of such hydrothermal salts, mostly composed of anhydrite, with lesser baryte, typically develop along oceanic seafloor ridges within heated subsurface fractures or at seafloor vents. There seawater-derived hydrothermal waters are heating, mixing, degassing, escaping and ultimately cooling. Active deep seafloor hydrothermal hydrologies create a specific group of sulphide ore deposits known as volcanic-hosted massive sulphide deposits (VHMS), with anhydrite as the primary-salt driving mineralisation.

The other non-solar salt grouping we shall discuss are salting-out precipitates, mostly halite, created when brines reach supercritical temperatures of 400-500°C. Some proponents of this mechanism postulate hydrothermal  halite sources much of the halite in active rifts such as the Red Sea or the Danakhil Depression (Hovland et al., 2006a, b).

Volcanogenic-hosted massive sulphide (VHMS) deposits

Volcanogenic-hosted massive sulphide deposits are forged by thermal circulation of seawater through newly-formed oceanic crust, in close temporal association with submarine volcanism. This milieu is characterised by active hydrothermal circulation and exhalation of metal sulphides, driven by mantle-induced geothermal gradients in oceanic basalt (Piercey et al., 2015). Being hosted in fractured basalts sets apart VHMS deposits from sedimentary exhalative (SedEx) and most sial-hosted Iron-Oxide-Copper-Gold (IOCG) deposits (Warren, 2016; Chapter 16). Hydrothermal anhydrite crystallises within a matrix of submarine volcanics and volcaniclastics via the heating of fissure-bound seawater (Figure 1a).

Anhydrite’s retrograde solubility across a range of salinities means the solubility of anhydrite decreases rapidly with increasing temperature in circulating seawater brines (Figure 1b; Blount and Dickson, 1969). Retrograde solubility also explains why anhydrite is most evident in the upper portions of vent mounds and in black and white “smokers.” Anhydrite's heating response is the opposite of baryte, another typical hydrothermal sulphate precipitate. Simple heating of seawater adjacent to seafloor vents, even without fluid mixing, will precipitate anhydrite, while simple cooling of hydrothermal waters will precipitate baryte. Once buried, hydrothermal calcium sulphate, in the presence of organic matter or hydrocarbons and circulating hydrothermal brines, acts as a sulphur source to create H2S, which then interacts with metal-carrying pore waters to co-precipitate metal sulphides.

Thus hydrothermal anhydrite, or more typically indicators of its former presence, are commonplace within volcanogenic hosted massive sulphide (VHMS) deposits. VHMS deposits usually form in submarine depressions as circulating seawater becomes an ore-forming hydrothermal fluid during interaction with the heated upper crustal rocks. Submarine depressions, especially those created by submarine calderas or by large-scale tectonic activity in median ocean-ridge rift valleys, are favourable sites and are often the home of an endemic chemosynthetic vent biota (Holden et al., 2012).


Based on VHMS Kuroko style deposits, fundamental processes typed to hydrothermal circulation and anhydrite distribution in a subduction zone as seen in the Koroko region of Japan, include (Ohmoto, 1996; Ogawa et al., 2007):


  • Intrusion of a heat source (typically ≈ 103 km size pluton) into oceanic crust or submarine continental crust causing deep convective circulation of seawater around the pluton (Figure 2). The radius of a typical circulation cell is ≈ 5 km. Temperatures of fluids discharging on to the seafloor increase with time from the ambient seafloor temperature to a typical maximum of ≈ 350° C, and then decrease gradually once more to ambient temperatures, on a time scale of ≈ 100 - 10,000 years. The majority of subsurface sulphide and sulphate mineralization occurs during the waxing stage of hydrothermal activity.

  • Reactions between warm country rocks and downward percolating seawater cause seawater SO4 to precipitate as disseminated and fracture-fill hydrothermal anhydrite in the country rock in areas where internal temperatures are greater than 150°C (Figures 2a, b: Sekko Anhydrite).
  • Reactions of “modified” seawater with higher-temperature rocks during the waning stages of hydrothermal circulation transform this sulphate-depleted “seawater” into metal-rich or H2S-rich ore-forming fluids. Metals are leached from the country rocks, while previously formed hydrothermal CaSO4 is reduced by Fe2+-bearing minerals and organic matter to provide H2S. The combined mass of high-temperature rocks that provide the metals and reduced sulphur in each VHMS marine deposit is typically ≈ 1011 tonnes (≈ 40 km3 in volume). Except for SO2, which produces acid-type alteration in some systems, the roles of magmatic fluids or gases are minor metal and sulphur in most massive sulphide systems.
  • Reactions between the ore-forming fluids and cooler rocks in the discharge zone cause zoned alteration of the rocks and precipitation of ore minerals in stockworkss.
  • Mixing of the ore-forming fluids with local seawater within unconsolidated sediments or on the seafloor can cause precipitation of “primitive ores” with a black ore mineralogy (Figure 2b; Oko Kuroko: sphalerite + galena + pyrite + baryte + anhydrite).
  • Reactions between the “primitive ores” with later and hotter hydrothermal fluids beneath a sulphate-rich thermal blanket cause a transformation of “primitive ores” to “matured ores” that are enriched in chalcopyrite and pyrite, often with a baryte cap in zones of cooling.

  • A mid-oceanic seafloor ridge region with significant documented volumes of anhydrite is located in the sediment-hosted Grimsey hydrothermal field in the Tjörnes fracture zone on the seafloor, north of Iceland (Figure 3a: Kuhn et al., 2003). There an active fracture zone is located at a ridge jump of 75 km, which caused widespread extension of the oceanic crust in this area. Hydrothermal activity in Grimsey field is spread over a 300 m by 1000 m area, at a water depth of 400 m. Active and inactive anhydrite chimneys up to 3 meters high and hydrothermal anhydrite mounds are typical of the seafloor in this area (Figure 3b-f). Clear, metal-depleted shimmering hydrothermal fluids, with temperatures up to 250°C, are venting from active chimneys and fluid inclusion in the precipitated anhydrites show the same homogenisation temperature range (Figure 3g).

    Anhydrite samples collected from Grimsey field average 21.6 wt.% Ca, 1475 ppm Sr and 3.47 wt.% Mg. The average molar Sr/Ca ratio is 3.3x10-3. Sulphur isotopes from anhydrites have typical δ34S seawater values of 22±0.7‰, indicating a seawater source for the SO4. Strontium isotopic ratios average 0.70662±0.00033, suggesting precipitation of anhydrite from a hydrothermal-seawater mixture (Figure 3f). The endmember of the venting hydrothermal fluids, calculated on a Mg-zero basis, contains 59.8 µmol/kg Sr, 13.2 mmol/kg Ca and a 87Sr/86Sr ratio of 0.70634. The average Sr/Ca partition coefficient between the hydrothermal fluids and anhydrite is about 0.67, implying precipitation from a non-evolved fluid. In combination, this suggests anhydrite forms in a zone of mixing between upwelling more deeply-seated hydrothermal fluids and shallowly circulating heated seawater (with a mixing ratio of 40:60). Before and during mixing, seawater is heated to 200-250°C, which drives anhydrite precipitation and the likely formation of an extensive anhydrite-rich zone beneath the seafloor, as in Hokuroko Basin.

    Once hydrothermal circulation slows or stops on a ridge or mound, and the “in-mound” temperature falls below 150°C, and anhydrite in that region tend to dissolve. During inactive periods, the dissolution leads to the collapse of sulphide chimneys and the internal dissolution of mound anhydrite. Additional ongoing disruption by faulting combine, so driving pervasive internal brecciation of the deposit. Through dissolution, former zones of hydrothermal anhydrite evolve into intervals of enhanced porosity and cavities in the mound. Such intervals initiate further fracture and collapse in the adjacent lithologies, which become permeable pathways during later renewed fluid circulation episodes. The alternating “coming and going” role of hydrothermal anhydrite creating precipitation space within the mound hydrology is similar to that of sedimentary evaporites in the sedimentary mineralising systems (Warren, 2016; Chapter 15).

    To form VHMS deposits on the seafloor, through-flushing hydrothermal fluids must transport sufficient amounts of metals and reduced sulphur, each at concentration levels > 1 ppm (Ohmoto, 1996). For a hydrothermal fluid with the salinity of normal seawater (≈0.7m ∑Cl) to be capable of transporting this amount of Cu and other base metals, it must be heated to temperatures > 300°C. Fluids with temperatures above 300°C will boil at pressures >200 bars. Under such conditions, the resulting vapour cannot carry sufficient quantities of metals to form a VHMS deposit. Boiling of a metalliferous hydrothermal brine outflow is prevented when the fluid vents into water that is deep enough to generate sufficient confining pressure. At 350°C, a minimum seawater depth of 1550m is necessary to prevent boiling. If the fluid passes through a sedimentary package where it loses temperature and metals (Cu, Ba) before emanating, the water depth beneath which boiling is prevented is less (≈1375m). Once vented, the turbulent mixing of hot hydrothermal waters with cooler seawater causes rapid precipitation of sulphides and calcium and barium sulphate, which produces the familiar black and white smokers (Blum and Puchelt, 1991).

    In modern oxic oceans, the sulphide-rich hydrothermal mounds are rapidly destroyed after the cessation of the hydrothermal activity (Herzig and Hannington, 1995; Tornos et al., 2015)). When hydrothermal activity at a mound decreases and the hydrothermal fluids cool to below 150 °C, the previously formed vent anhydrite is dissolved (retrograde solubility). This near-surface cooling contributes to the dissolution collapse of the anhydrite supported mound surface, particularly at the mound flanks, and allows additional influx of cold seawater. As mound flank collapse expands the remaining detrital pyritic sand residues are replaced by oxyhydroxides, and copper sulphides tend to be oxidised and replaced by atacamite (Knott et al., 1998). If seafloor weathering continues to completion, all the metal sulphides become oxidized or dissolved. Only those metalliferous VHMS deposits capped by impermeable volcanic, volcaniclastic, or sedimentary deposits soon after formation are preserved due to shielding from the oxidising conditions at the deep seafloor.

    In all cases, VHMS deposit styles of mineralisation, along with associated anhydrite precipitation, are allied to submarine volcanism and hydrothermally-driven circulation of seawater within adjacent deepwater sediments. Hydrothermal anhydrite typifies mineralisation in a variety of tectonic settings and sediment types (more detail on deposit styles and their anhydrites are given in Warren, 2016; Chapter 16):

  • VHMS deposits formed in subduction-related island-arc settings (Kuroko-type deposits; Ogawa et al., 2007);
  • VHMS deposits formed at mid-oceanic or back-arc spreading centres (Grimsey vent field or TAG mound type deposits; Kuhn et al., 2003; Knott et al., 1998);
  • VHMS deposits formed at spreading centres, but, due to the proximity of one or more landmasses, the deposit is sediment-hosted (Besshi-type deposits). This style of deposit shows some affinities with SedEx deposits but, unlike a SedEx deposit, the hydrological drive is linked to igneous intrusion.
  • In all cases, vestiges of the once voluminous anhydrite are a minor component in the cooled brecciated and fractured volcanic pile. This variety of CaSO4, and its variably metalliferous pseudomorphs and breccias, are not associated with solar heating. The occurrence sparry hydrothermal anhydrite as fracture and breccia fill in a labile volcanic pile, always covered with deep ocean sediments, cherts, etc., makes the distinction from sedimentary anhydrite relatively straightforward.

    Halite and a hydrothermal brine's critical temperature

    In terms of a significant volume of non-evaporite salt produced, or the possible ubiquity of a non-solar process contributing large amounts of NaCl, is the possible formation of halite when a brine reaches supercritical temperatures at appropriate depths in tectonically active parts of the earth's crust. Starting in the mid-2000's Hovland et al. (2006a,b) then Hovland and Rueslatten (2009) introduced the concept of substantial volumes of hydrothermal halite precipitating from subsurface brines at supercritical temperatures, especially in the buried hot portions of thermally-active rift basement. Two recent papers Hovland et al. (2018a,b) summarise much of this earlier material and add the notion of serpentinization being sink for chloride and  a driver of halite formation in many evaporite basins. Countering arguments to the notion of a non-evaporite origin for substantial volumes of halite in sedimentary basins are given in Talbot (2008) and Aftabi and Atapour (2018). The notion of the importance of the Wilson cycle in a sedimentary evaporite (megahalite and megasulphate basins) context, rather than a direct igneous-metamorphic source as argued by Hovland, is summarised in Warren (2016, Chapter 5).

    A Hovland model of a non-evaporite source of halite relies on heated subsurface brines becoming supercritical and so transforming a brine to a fluid that does not dissolve but precipitates salt (within specific temperature and pressure ranges). A supercritical fluid is defined as any substance at a temperature and pressure above its critical point; in such a state, it can effuse through solids like a gas, and dissolve materials like a liquid. In addition, close to the critical point, small changes in pressure or temperature result in substantial changes in density. The critical point (CP), also called a critical state, specifies the conditions (temperature, pressure and sometimes composition) at which a phase boundary ceases to exist. At particular pressure/temperature conditions, supercritical water is unable to dissolve/retain common sea salts in solution (Josephson, 1982; Bischoff and Pitzer, 1989; Simoneit 1994; Hovland et al., 2006a).

    When seawater brines are heated in pressure cells in the laboratory, they pass into the supercritical region at a temperature of 405°C and 300 bar pressure (the CP of seawater). A particulate ‘cloud’ then forms via the onset of ‘shock crystallization’ of NaCl and Na2SO4 (Figure 4a). The sudden phase transition occurs as the solubility of the previously dissolved salts declines to near-zero, across a temperature range of only a few degrees, and is associated with a substantial lowering of density (Figure 4b). The resulting solids in the “cloud” consist of amorphous microscopic NaCl and Na2SOparticles with sizes between 10 and 100 mm. The resultant “salting out” can lead to the precipitation of large volumes of subsurface salts in fractures and fissures and perhaps even in the deeper portions of salt structures. The same supercritical conditions improve the ability of brines to carry high volumes of hydrothermal hydrocarbons prior to the onset of supercritical conditions (Josephson, 1982; McDermott et al., 2018). Supercritical water has enhanced solvent capacity for organic compounds and reduced solvation properties for ionic species due to its loss of aqueous hydrogen bonding (Figure 4c; Simoneit, 1994).

    Hovland et al. (2006a, b) predict that some of the large volumes of deep subsurface salt found in the Red Sea, in the Mediterranean Sea and the Danakil depression, formed via the forced magmatically-driven hydrothermal circulation of seawater down to depths where it became supercritical. This salt, they argue, was precipitated deep under-ground via “shock crystallisation” from a supercritical effusive phase and so formed massive accumulations (mostly halite) typically in crustal fractures that facilitated the deep circulation. NaCl then flowed upwards in solution in dense, hot hydrothermal brine plumes, precipitating more solid salt beds upon cooling nearer or on the surface/seafloor. More recently, Scribano et al. (2017) and Hovland et al. (2018a, b) have added the argument that serpentinisation is the dominant source of halite in the Messinian succession of the Mediterranean.

    To date, the Hovland et al. model of hydrothermal sourcing for widespread halite from a supercritical brine source (in active magmatic settings) has not been widely accepted by the geological community (Talbot, 2008; Warren., 2016; Aftabi and Atapour, 2018). To date, no direct indications of the formation of masses of halite formed by this process have been sampled. In contrast to the theories of Hovland et al (2018b), textures in the potash and halite salts in the Danakhil depression are evaporitic with only small volumes of hydrothermal overprint driven by the escape of saline volatiles derived thermal decomposition of hydrated salts. The postulated diapiric structures are not present in seismic, nor are any other buried hydrothermal/halokinetic structures visible in seismic (Bastow et al., 2018; Salty Matters; Warren 2016). Likewise, all the features seen in core and seismic in the Messinian of the Mediterranean are layered with classic sedimentary and halokinetic textures. The seismic across the Red Sea salt structures and the layering in the brine deeps are easily explained by current sedimentary and layered deep seafloor ponded brine (DHAL) models.

    The high temperatures required for supercritical seawater venting mean such sites are rare on the seafloor. The deepest thermal upwelling site where supercritical seafloor conditions are thought to be active just below the upwelling site is the Beebe vent field (Figure 5; Webber et al., 2015; McDermott et al., 2018). At 4960 m below sea level, the vent field sits atop the ultra-slow spreading Mid Cayman Rise and is the world’s deepest known hydrothermal exhalative system. Situated on very thin (2–3 km thick) oceanic crust at an ultraslow spreading centre, this hydrothermal system circulates fluids to depths ≈1.8 km in a basement that is likely to include a mixture of both mafic and ultramafic lithologies (Webber et al., 2015).

    The surface of the active vent field is made up of high temperature (≈401°C) anhydritic ‘‘black smokers’’ that build Cu, Zn and Au-rich sulfide mounds and chimneys (Figure 5a). The vent field is highly gold-rich, with Au values up to 93 pp, with an average Au:Ag ratio of 0.15. Gold precipitation is directly associated with diffuse flow through anhydritic‘‘beehive’’ chimneys. Significant mass-wasting of sulfide material in the vent field, accompanied by changes in metal content results in metalliferous talus and interfingering with deep marine sediment deposits (Figure 5b, c, d, e).

    All the high-temperature endmember fluids venting at the Beebe site show Cl levels that are significantly lower than seawater, with an average endmember concentration of 349 mmol/kg. Due to the lack of a significant sink for Cl within mafic-hosted subsurface circulation pathways, Cl depletions in vent fluids are typically attributed to phase separation (McDermott et al., 2018). Thus, the intrinsic Cl depletions, in conjunction with a seafloor pressure of 496 bar, places the two-phase boundary at 483°C, suggesting that escaping fluids experienced a phase separation at conditions that are both hotter and deeper (higher pressure) than the critical point for seawater at 407°C and 298 bar (Bischoff, 1991).

    During incipient phase separation from supercritical seawater, a small amount of high-salinity brine it thought to condense in the subsurface as a separate phase, so creating the Cl-depleted residual fluid, or vapour phase. The Cl-depleted fluids venting at Beebe are thought to represent this vapour phase. Although a vent fluid of seawater chlorinity is not a supercritical fluid at the conditions of seafloor venting (398°C, 496 bar), the vent fluids indicate sourcing from a supercritical phase owing to their lower chlorinity (Bischoff and Pitzer, 1989).

    Phase relations in the system NaCl-H2O (Bischoff, 1991) can be used to estimate the minimum temperature of phase separation at the Beebe site, based on the chlorinity of the vapour phase and the assumption that phase separation occurs at or below the seafloor. A minimum temperature of 491°C is required to produce the measured Cl concentration of 349 mmol/kg observed in the Beebe Vents endmember fluids. Accordingly, the observed Cl depletion in the high-temperature endmember fluids implies that these fluids must have cooled by at least 90°C prior to venting, and perhaps more (McDermott et al., 2018). For example, if the location of phase separation was 1000 m deeper, then that would require maximum fluid temperatures in the vicinity of separation of 535°C.

    The only salt that can be confused with an evaporite salt at the Beebe site is hydrothermal anhydrite. Any halite or Na2SOderived from seawater reaching its supercritical point is still located in fissures many hundreds of metres below the surface and is as yet unsampled. Likewise, there are no halite-saturated brine ponds on the seafloor and the smoker anhydrite, like much of the metalliferous content, has a low preservation potential as it is being leached back into seawater via galvanic interaction (Webber et al., 2015).

    Herein is the problem for assessing the viability of a Hovlnd-style model for halite. Where is the evidence and the data? Until significant volumes of hydrothermal halite are intersected somewhere on the earth's surface, there is not a working example, only a sophisticated reinterpretation of existing halite occurrences. Modern seawater (rather than an experimental NaCl -H2O system) would give not just halite but also Na2SOat its critical point, where are the volumes or texural and mineralogical indications of these salts, or their brines and alteration haloes? Until there is the physical proof of a working example of substantial hydrothermal halite sourced in supercritical phase separations, I prefer to apply Occam's Razor.

    Halite alteration, renewed deep brine flow and metamorphism

    A source of chlorine-rich hydrothermal fluid (not halite) in the deep subsurface is the recycling of deeply buried sedimentary mega-halite units into the greenschist realm and beyond (Yardley and Graham, 2002). In the metamorphic realm (T>200°C) the derived fluids do not precipitate halite, but a series of meta-evaporite indicator minerals (Table 1). Lewis and Holness (1996) demonstrated that buried salt bodies, subjected to high pressures and elevated temperatures, can acquire a permeability comparable to that of a sand, within what is sometimes called the "Holness zone". This is because the crystalline structure of deeply buried salt (halite) attains dihedral angles between salt crystals of less than 60 degrees, and so creates an impermeable polyhedral meshwork (Figure 6). Such conditions probably begin at the onset of greenschist P-T conditions, whereby highly-saline hot brines form continuous brine stringers around all such altered and recrystallizing salt crystals.

    This polyhedral permeability meshwork allows hot dense brines or hydrocarbons to migrate through salt (Schoenherr et al., 2007a, b) and ultimately dissolve the salt host, releasing a pulse of sodic- and chloride-rich fluid into the metamorphic realm (Warren, 2016; Chapter 14). It is why little or no evidence of solid masses of metamorphosed halite is found in subsurface meta-evaporitic settings where temperatures have exceeded 250 - 300 °C, even though the melting point of halite is 800°C. Given the right subsurface conditions these halite-derived metamorphic brines may evolve into supercritical waters.

    Contrary to conventional geological modelling of salt in diapirs being mostly  impermeable, Hovland et al. (2018a) argue for the formation of salt stocks by hot brines migrating upward through the middle of the salt body; provided that the salt stock is situated within the "Holness zone." This assumes that "Holness zone" flows brine and can also reach subcritical conditions. The inferred rising flow of intrasalt hot brines then reach saturation upon cooling in the upper part of the salt stem, where solid salts are precipitating according to their specific solubility at each particular temperature and pressure interval. The Hovland model thus includes a refining process in the salt stem, where halite, for example, precipitates upon cooling long before calcium and magnesium chloride salts. However, in the Danakhil Depression, where they infer this process is active (Hovland et al., 2018a), the seismic indicates the evaporite mass is bedded and faulted, while the evaporite textures recovered in cores and doline/uplift landforms across the saltflat surface combine to show Holness-zone halokinesis is not segregating the halite, kainite, carnallite, sylvite and bischofite salts that typify the region around the Dallol Mound (Bastow et al., 2018; Warren, 2016, Chapter 11).

    Beyond the greenschist facies and the polyhedral transition of sedimentary/halokinetic halite, metamorphic minerals with an evaporite protolith tend to be enriched in minerals entraining sodium, potassium and magnesium (Figure 7; Table 1; Yardley and Graham, 2002). These metamorphic minerals (meta-evaporites) can entrain high levels of volatiles (Cl, SO3 and CO2) as well as elevated levels of boron, along with high salinity in the associated metamorphic fluids; all indicate their evaporitic protolith (Table 1; Figure 7).

    Sodium tends to come from the dissolution of salts, such as halite, kainite or trona; while magnesium tends to be remobilized from earlier diagenetic minerals, such as reflux dolomites,magnesium-rich evaporitic clays and some potash minerals (Table 1). Boron in tourmalinites may have come from a colemanite/ulexite lacustrine precursor. Once direct evidence of a salty protolith is largely removed via fluid dispersion in burial and ongoing loss of volatiles, the palaeo-evaporite indications are restricted mostly to mineralogic associations, along with an occasional textural relict of a former evaporitic breccia bed, rauwacke or salt weld (Warren, 2016; Chapters 7, 14).

    Evidence of early stages in an evaporite-fed sodic transformation is seen in the sodic phlogopites (phlogopite = magnesian mica) and sodian aluminian talcs in the metapelites of the Tell Atlas in Algeria (Schreyer et al. 1980). Evaporitic sulphate crystals are pseudomorphed in the NaCl-scapolite-dominated sequences of the Cordilleras Beticas of Spain (Gómez-Pugnaire et al. 1994). Rocks of higher temperature and pressure facies, such as the massive stratiform anorthosites in the Grenville Precambrian Province of North America, have been interpreted as possible meta-evaporites (Gresens, 1978), as have the anhydrite-containing Mesoproterozoic calcsilicates in the Oaxacan granulite complex in southern Mexico (Ortega-Gutierrez, 1984) and the pervasive scapolites in the Neoproterozoic Zambesian orogenic belt of Zambia (Hanson et al., 1994). Subsequent work on the Grenville anorthosites, although still allowing for a metasedimentary protolith, has concluded an igneous source of volatiles is more likely (Moecher et al., 1992; Peck and Valley, 2000; Glassly et al., 2010). Defining the likelihood of an evaporite protolith becomes increasingly difficult as the metamorphic grade increases. Once a metamorphic rock enters the granulite facies, its protolith interpretation is typically much more contentious (e.g. the evaporite versus carbonatite interpretations in the Oaxacan granulites, Mexico).

    Hydrothermal gypsum

    Some of the most visually striking examples of hydrothermal gypsum precipitation are in the Naica mine, Chihuahua, Mexico (Figure 8). There several natural caverns, such as Cave of Swords (Cueva de la Espades discovered in 1975) and Cave of Crystals (Cueva de los Cristales discovered in 2000), contain giant, faceted, and transparent single crystals of gypsum as long as 11 m (Figure 9a; García-Ruiz et al., 2007; Garofalo et al., 2010). Crystals in Cueva de los Cristales are the largest documented gypsum crystals in the world. These huge crystals grew slowly at very low supersaturation levels from thermal phreatic waters with temperatures near the gypsum-anhydrite boundary. Gypsum still precipitates today on mine walls.

    According to García-Ruiz et al., 2007, the sulphur and oxygen isotopic compositions of these gypsum crystals are compatible with growth from solutions resulting from dissolution of anhydrite, which was previously precipitated during late hydrothermal mineralisation in a volcanogenic matrix. The chemistry suggests that these megacrystals formed via a self-feeding mechanism, driven by a solution-mediated, anhydrite-gypsum phase transition. Nucleation kinetics calculations based on laboratory data show that this mechanism can account for the formation of these giant crystals, yet only when operating within a very narrow range of temperature of a few degrees as identified by the fluid inclusion values.

    Fluid inclusion analyses show that the giant crystals came from low-salinity solutions at temperatures ≈ 54°C, slightly below the temperature of 58°C where the solubility of anhydrite equals that of gypsum (Figure 9b; García-Ruiz et al., 2007). Van Driessche et al. (2011) argue the slowest gypsum crystal growth in the phreatic cavern occurred when waters were at 55°C. At this temperature, the crystals would take 990,000 years to grow to a diameter of 1 meter. By increasing the temperature in the cave by one degree, to 56° C, the same size crystal could have formed in a little less than half the time, or around 500,000 years. This possible growth rate would work out to about a billionth of a meter of growth per day and is perhaps the slowest growth rate that has ever been measured.

    Garofolo et al., 2010, accept the need for a limited temperature range during precipitation, but argue the precipitating solutions were in part meteorically influenced. Their work focused on Cueva de las Espadas. As for most other hypogenic caves, prior to their analytical work, they assumed that caves of the Naica region lacked a direct connection with the land surface and so gypsum precipitation would be unrelated to climate variation. Yet, utilising a combination of fluid inclusion and pollen spectra data from cave and mine gypsum, they concluded climatic changes occurring at Naica exerted and influence on fluid composition in the Espadas caves, and hence on crystal precipitation and growth.

    Microthermometry and LA-ICP-Mass Spectrometry of fluid inclusions in the gypsum in the Cueva de las Espadas indicate that brine source was a shallow, chemically peculiar, saline fluid (up to 7.7 eq. wt.%NaCl) and that it may have formed via evaporation, during an earlier dry and hot climatic period. In contrast, the fluid of the deeper caves (Cristales) was of lower salinity (≈3.5 eq. wt.% NaCl) and chemically homogeneous, and likely was little affected by evaporation processes. Galofolo et al. (2010) propose that mixing of these two fluids, generated at different depths of the Naica drainage basin, determined the stable supersaturation conditions needed for the gigantic gypsum crystals to grow (Figure 9c). The hydraulic communication between Cueva de las Espadas and the other deep Naica caves controlled fluid mixing. Mixing must have taken place during alternating cycles of warm-dry and fresh-wet climatic periods, which are known to have occurred in the region. Pollen grains from 35 ka-old gypsum crystals from the Cave of Crystals indicates a relatively homogenous catchment basin dominated by a mixed broadleaf wet forest. This suggests precipitation during a fresh-wet climatic period; the debate continues as to whether the gypsum at Naica is a mixing zone or a hydrothermal salt.

    Solar versus nonsolar salts

    This and the previous article show that substantial volumes of various salts (especially retrograde anhydrite) form in the terrestrial subsurface, independent of solar evaporation. Except for some bedded cryogenic salt bodies (e.g., Korabogazgol in Kazakhstan or Noachian lakes on Mars), non-solar evaporation salts tend to nucleate in subsurface fractures, and breccia interspaces in the igneous and metamorphic realm. Crystals  tend to be cavity cements, but can also replace portions of a pre-existing salt mass. On Earth, the most widespread non-solar salt is anhydrite with occurrences ranging from volcanic hosted mid-ocean ridges to Kuroko style deposits in subduction zones. In all cases, the intimate association with submarine volcanics and fissures, where hydrothermally heated seawater once circulated, mean this type of hydrothermal (non solar heating) salt is readily distinguished from sedimentary anhydrite.

    For halite, there is little direct evidence of any massive halite occurrence in outside of sedimentary basins where isolated-sumps of ponded brine were once evaporated. A sophisticated notion theorising hydrothermal halite has been published by Hovland and co-workers (e.g. Hovland et al., 2018a b; Scribano et al., 2017) to explain some halite occurrences in tectonically-active areas. There is little support for this model in the published literature outside of Hovland and co-workers. Thick buried solar halite masses tend form in particular stages of the Wilson Cycle. Once buried, these evaporite masses are mostly impervious, they can flow and dissolve, and on entry into the greenschist realm can become permeable, so feeding large volumes of highly saline brines into the metamorphic and igneous realms. These brines can drive metal accumulations and the formation of characteristic meta-evaporitic minerals and gemstones (Warren, 2016).


    Aftabi, A., and H. Atapour, 2018, Comment on the papers by Hovland et al., 2018b, Hovland et al., 2018a “Large salt accumulations as a consequence of hydrothermal processes associated with 'Wilson cycles': A review” (part 1 and 2): Marine and Petroleum Geology, v. 98, p. 890-897.

    Ayuso, R. A., and C. E. Brown, 1984, Manganese-rich red tourmaline from the Fowler talc belt, New York: Canadian Mineralogist, v. 22, p. 327-331.

    Bastow, I. D., A. D. Booth, G. Corti, D. Keir, C. Magee, C. A.-L. Jackson, J. Warren, J. Wilkinson, and M. Lascialfari, 2018, The development of late-stage continental breakup: Seismic reflection and borehole evidence from the Danakil Depression, Ethiopia: Tectonics, v. 37.

    Bischoff, J. L., and K. S. Pitzer, 1989, Liquid-vapor relations for the system NaCl-H2O; summary of the P-T-x surface from 300 degrees to 500° C: American Journal of Science, v. 289, p. 217-248.

    Bischoff, R., 1991, Lithological interpretation of a gas cavern well in Zechstein evaporites of the Lesum salt dome (Bremen, Germany): Oil Gas European Magazine, v. 17, p. 16-19.

    Blount, C. W., and F. W. Dickson, 1969, The solubility of anhydrite (CaSO4) in NaCl-H2O from 100 to 450° C and 1 to 1000 bars: Geochimica et Cosmochimica Acta, v. 33, p. 227-245.

    Blum, N., and H. Puchelt, 1991, Sedimentary-hosted polymetallic massive sulphide deposits of the Kebrit and Shaban Deeps, Red Sea.: Mineralium Deposita, v. 26, p. 217-227.

    Byerly, G. R., and M. R. Palmer, 1991, Tourmaline mineralization in the Barberton greenstone belt, South Africa; early Archean metasomatism by evaporite-derived boron: Contributions to Mineralogy and Petrology, v. 107, p. 387-402.

    Cook, N. D. J., and P. M. Ashley, 1992, Meta-evaporite sequence, exhalative chemical sediments and associated rocks in the Proterozoic Willyama Supergroup, South Australia: implications for metallogenesis: Precambrian Research, v. 56, p. 211-226.

    Davidson, G. J., and R. R. Large, 1994, Gold metallogeny and the copper-gold association of the Australian Proterozoic: Mineralium Deposita, v. 29, p. 208 - 223.

    Dostal, J., J. Keppie, H. Macdonald, and F. Ortega-Gutiérrez, 2004, Sedimentary Origin of Calcareous Intrusions in the ~1 Ga Oaxacan Complex, Southern Mexico: Tectonic Implications: International Geology Review, v. 46, p. 528-541.

    Eglinger, A., C. Ferraina, A. Tarantola, A.-S. André-Mayer, O. Vanderhaeghe, M.-C. Boiron, J. Dubessy, A. Richard, and M. Brouand, 2014, Hypersaline fluids generated by high-grade metamorphism of evaporites: fluid inclusion study of uranium occurrences in the Western Zambian Copperbelt: Contributions to Mineralogy and Petrology, v. 167, p. 1-28.

    Engvik, A. K., K. Mezger, S. Wortelkamp, R. Bast, F. Corfu, A. Korneliussen, P. Ihlen, B. Bingen, and H. Austrheim, 2011, Metasomatism of gabbro – mineral replacement and element mobilization during the Sveconorwegian metamorphic event: Journal Of Metamorphic Geology, v. 29, p. 399-423.

    Evans, D. M., 2017, Fe-Ni-Cu sulfide-evaporite association at Munali, Zambia: Society of Geology Applied to Mineral Deposits Biennial Meeting, Québec City, Province de Québec.

    Evans, R., 1970, Genesis of sylvite- and carnallite-bearing rocks from Wallace, Nova Scotia: Third Symposium on Salt, v. 1, p. 239-245.

    Faryad, S. W., 2002, Metamorphic Conditions and Fluid Compositions of Scapolite-Bearing Rocks from the Lapis Lazuli Deposit at Sare Sang, Afghanistan: Journal of Petrology, v. 43, p. 725-747.

    Feneyrol, J., G. Giuliani, D. Ohnenstetter, A. E. Fallick, J. E. Martelat, P. Monié, J. Dubessy, C. Rollion-Bard, E. Le Goff, E. Malisa, A. F. M. Rakotondrazafy, V. Pardieu, T. Kahn, D. Ichang'i, E. Venance, N. R. Voarintsoa, M. M. Ranatsenho, C. Simonet, E. Omito, C. Nyamai, and M. Saul, 2013, New aspects and perspectives on tsavorite deposits: Ore Geology Reviews, v. 53, p. 1-25.

    Feneyrol, J., D. Ohnenstetter, G. Giuliani, A. E. Fallick, C. Rollion-Bard, J.-L. Robert, and E. P. Malisa, 2012, Evidence of evaporites in the genesis of the vanadian grossular tsavorite deposit in Namalulu, Tanzania: The Canadian Mineralogist, v. 50, p. 745-769.

    Frietsch, R., P. Tuisku, O. Martinsson, and J. Perdahl, 1997, Early Proterozoic Cu-(Au) and Fe ore deposits associated with regional NaCl metasomatism in northern Fennoscandinavia: Ore Geology Reviews, v. 12, p. 1-34.

    García-Ruiz, J. M., R. Villasuso, C. Ayora, A. Canals, and F. Otálora, 2007, Formation of natural gypsum megacrystals in Naica, Mexico: Geology, v. 35, p. 327-330.

    Garnier, V., G. Giuliani, D. Ohnenstetter, A. E. Fallick, J. Dubessy, D. Banks, H. Q. Vinh, T. Lhomme, H. Maluski, A. Pecher, K. A. Bakhsh, P. Van Long, P. T. Trinh, and D. Schwarz, 2008, Marble-hosted ruby deposits from Central and Southeast Asia: Towards a new genetic model: Ore Geology Reviews, v. 34, p. 169-191.

    Garofalo, P. S., M. B. Fricker, D. Günther, P. Forti, A.-M. Mercuri, M. Loreti, and B. Capaccioni, 2010, Climatic control on the growth of gigantic gypsum crystals within hypogenic caves (Naica mine, Mexico)?: Earth and Planetary Science Letters, v. 289, p. 560-569.

    Giuliani, G., A. Cheilletz, C. Arboleda, V. Carrillo, F. Rueda, and J. H. Baker, 1995, An evaporitic origin of the parent brines of Colombian emeralds; fluid inclusion and sulphur isotope evidence: European Journal of Mineralogy, v. 7, p. 151-165.

    Giuliani, G., J. Dubessy, D. Ohnenstetter, D. Banks, Y. Branquet, J. Feneyrol, A. E. Fallick, and J.-E. Martelat, 2017, The role of evaporites in the formation of gems during metamorphism of carbonate platforms: a review: Mineralium Deposita, p. 1-20.

    Giuliani, G., G. R. Olivo, O. J. Marini, and D. Michel, 1993, The Santa Rita gold deposit in the Proterozoic Paranoa Group, Goias, Brazil; an example of fluid mixing during ore deposition: Ore Geology Reviews, v. 8, p. 503-523.

    Glassley, W. E., J. A. Korstgård, and K. Sørensen, 2010, K-rich brine and chemical modification of the crust during continent-continent collision, Nagssugtoqidian Orogen, West Greenland: Precambrian Research, v. 180, p. 47-62.

    Gómez-Pugnaire, M. T., G. Franz, and V. L. Sanchez-Vizcaino, 1994, Retrograde formation of NaCl-scapolite in high pressure metaevaporites from the Cordilleras-Beticas (Spain): Contributions to Mineralogy & Petrology, v. 116, p. 448-461.

    Gresens, R. L., 1978, Evaporites as precursors of massif anorthosite: Geology, v. 6, p. 46-50.

    Hanson, R. E., T. J. Wilson, and H. Munyanyiwa, 1994, Geologic evolution of the Neoproterozoic Zambesi Orogenic belt in Zambia [Review]: Journal of African Earth Sciences & the Middle East., v. 18, p. 135-150.

    Heimann, A., P. G. Spry, G. S. Teale, W. R. Leyh, C. H. H. Conor, G. Mora, and J. J. O'Brien, 2013, Geochemistry and genesis of low-grade metasediment-hosted Zn-Pb-Ag mineralization, southern Proterozoic Curnamona Province, Australia: Journal of Geochemical Exploration, v. 128, p. 97-116.

    Herzig, P. M., and M. D. Hannington, 1995, Polymetallic massive sulphides at the modern sea floor: Ore Geology Reviews, v. 10, p. 95-115.

    Hogarth, D. D., 1979, Lapis lazuli from Edwards, New York; a possible metaevaporite: Geol. Assoc. Can. Mineral. Assoc. Can., Jt. Annu. Meet., Program Abstr.

    Hogarth, D. D., and W. L. Griffin, 1978, Lapis lazuli from Baffin Island; a Precambrian meta-evaporite: Lithos, v. 11, p. 37-60.

    Holden, J. F., J. A. Breier, K. L. Rogers, M. D. Schulte, and B. M. Toner, 2012, Biogeochemical Processes at Hydrothermal Vents

    Microbes and Minerals, Bioenergetics, and Carbon Fluxes: Oceanography, v. 25, p. 196-208.

    Hovland, M., H. K. Johnsen, and H. Rueslåtten, 2019, Salt-formation in rifting and subduction (Wilson cycles): Reply to Alijan Aftabi and Habibeh Atapour on their comments to our two articles: Marine and Petroleum Geology, v. 100, p. 554-558.

    Hovland, M., T. Kuznetsova, H. Rueslatten, B. Kvamme, H. K. Johnsen, G. E. Fladmark, and A. Hebach, 2006a, Sub-surface precipitation of salts in supercritical seawater: Basin Research, v. 18, p. 221-230.

    Hovland, M., and H. Rueslatten, 2009, Origin and permeability of deep ocean salts: Geophysical Research Abstracts, v. 11.

    Hovland, M., and H. Rueslatten, 2009, Origin and permeability of deep ocean salts: Geophysical Research Abstracts, v. 11.

    Hovland, M., H. Rueslåtten, and H. K. Johnsen, 2018a, Large salt accumulations as a consequence of hydrothermal processes associated with ‘Wilson cycles’: A review Part 1: Towards a new understanding: Marine and Petroleum Geology, v. 92, p. 987-1009.

    Hovland, M., H. Rueslåtten, and H. K. Johnsen, 2018b, Large salt accumulations as a consequence of hydrothermal processes associated with ‘Wilson cycles’: A review, Part 2: Application of a new salt-forming model on selected cases: Marine and Petroleum Geology, v. 92, p. 128-148.

    Hovland, M., H. G. Rueslatten, H. K. Johnsen, B. Kvamme, and T. Kuznetsova, 2006b, Salt formation associated with sub-surface boiling and supercritical water: Marine and Petroleum Geology, v. 23, p. 855-869.

    Hunt, J. A., 2005, The geology and genesis of iron oxide copper-gold mineralisation associated with Wernecke Breccia, Yukon, Canada: Doctoral thesis, James Cook University, Townsville, Australia, 120 p.

    Josephson, J., 1982, Supercritical fluids: Environmental Science and Technology, v. 16, p. 548A-551A.

    Kendrick, M. A., M. Honda, D. Gillen, T. Baker, and D. Phillips, 2008, New constraints on regional brecciation in the Wernecke Mountains, Canada, from He, Ne, Ar, Kr, Xe, Cl, Br and I in fluid inclusions: Chemical Geology, v. 255, p. 33-46.

    Knott, R., Y. Fouquet, J. Honnorez, S. Petersen, and M. Bohn, 1998, 1. Petrology of hydrothermal mineralization: A vertical section through TAG Mound, in P. M. Herzig, S. E. Humphris, D. J. Miller, and R. A. Zierenberg, eds., Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 158.

    Kuhn, T., P. M. Herzig, M. D. Hannington, D. Garbe-Schönberg, and P. Stoffers, 2003, Origin of fluids and anhydrite precipitation in the sediment-hosted Grimsey hydrothermal field north of Iceland: Chemical Geology, v. 202, p. 5-21.

    Kukla, P., J. Urai, J. K. Warren, L. Reuning, S. Becker, J. Schoenherr, M. Mohr, H. van Gent, S. Abe, S. Li, Desbois, G. Zsolt Schléder, and M. de Keijzer, 2011a, An Integrated, Multi-scale Approach to Salt Dynamics and Internal Dynamics of Salt Structures: AAPG Search and Discovery Article #40703 (2011).

    Lewis, S., and M. Holness, 1996, Equilibrium halite-H2O dihedral angles: High rock salt permeability in the shallow crust: Geology, v. 24, p. 431-434.

    Markl, G., and S. Piazolo, 1998, Halogen-bearing minerals in syenites and high-grade marbles of Dronning Maud Land, Antarctica - Monitors of fluid compositional changes during late-magmatic fluid-rock interaction processes: Contributions to Mineralogy & Petrology, v. 132, p. 246-248.

    McDermott, J. M., S. P. Sylva, S. Ono, C. R. German, and J. S. Seewald, 2018, Geochemistry of fluids from Earth’s deepest ridge-crest hot-springs: Piccard hydrothermal field, Mid-Cayman Rise: Geochimica et Cosmochimica Acta, v. 228, p. 95-118.

    Moecher, D. P., E. J. Essene, and J. W. Valley, 1992, Stable isotopic and petrological constraints on scapolitization of the Whitestone meta-anorthosite, Grenville Province, Ontario: Journal Of Metamorphic Geology, v. 10, p. 745-762.

    Morteani, G., Y. A. Kostitsyn, H. A. Gilg, C. Preinfalk, and T. Razakamanana, 2013, Geochemistry of phlogopite, diopside, calcite, anhydrite and apatite pegmatites and syenites of southern Madagascar: evidence for crustal silicocarbonatitic (CSC) melt formation in a Panafrican collisional tectonic setting: International Journal of Earth Sciences, v. 102, p. 627-645.

    Ogawa, Y., N. Shikazono, D. Ishiyama, H. Sato, T. Mizuta, and T. Nakano, 2007, Mechanisms for anhydrite and gypsum formation in the Kuroko massive sulfide–sulfate deposits, north Japan: Mineralium Deposita, v. 42, p. 219-233.

    Ohmoto, H., 1996, Formation of volcanogenic massive sulphide deposits - The Kuroko perspective: Ore Geology Reviews, v. 10, p. 135-177.

    Oliver, N. H. S., 1995, Hydrothermal history of the Mary Kathleen Fold belt, Mt Isa Block, Queensland: Australian Journal of Earth Sciences, v. 42, p. 267-279.

    Ortega-Gutierrez, F., 1984, Evidence of Precambrian evaporites in the Oaxacan granulite complex of southern Mexico: Precambrian Research, v. 23, p. 377-393.

    Parente, C., L. Ronchi, A. Sial, J. Guillou, M. Arthaud, K. Fuzikawa, and C. Veríssimo, Ceara, 2004, Geology and geochemistry of Paleoproterozoic magnesite deposits (≈1.8Ga), State of Ceará, Northeastern Brazil: Carbonates and Evaporites, v. 19, p. 28-50.

    Peck, W. H., and J. W. Valley, 2000, Large crustal input to high δ18O anorthosite massifs of the southern Grenville Province: new evidence from the Morin Complex, Quebec: Contributions to Mineralogy and Petrology, v. 139, p. 402-417.

    Peng, Q. M., and M. R. Palmer, 1995, The Palaeoproterozoic boron deposits in eastern Liaoning, China: a metamorphosed evaporite: Precambrian Research, v. 72, p. 185-197.

    Piercey, S. J., J. M. Peter, and R. J. Herrington, 2015, Zn-rich Volcanogenic Massive Sulphide (VMS) Deposits: Current Perspectives on Zinc Deposits, Irish Association for Economic Geology, 37-57 p.

    Plimer, I. R., 1988, Tourmalinites associated with Australian Proterozoic submarine exhalative ores, in G. H. Friedrich, and P. M. Herzig, eds., Base Metal Sulphide Deposits: Berlin, Springer Verlag, p. 255-283.

    Schoenherr, J., R. Littke, J. L. Urai, P. A. Kukla, and Z. Rawahi, 2007a, Polyphase thermal evolution in the Infra-Cambrian Ara Group (South Oman Salt Basin) as deduced by maturity of solid reservoir bitumen: Organic Geochemistry, v. 38, p. 1293-1318.

    Schoenherr, J., J. L. Urai, P. A. Kukla, R. Littke, Z. Schleder, J.-M. Larroque, M. J. Newall, N. Al-Abry, H. A. Al-Siyabi, and Z. Rawahi, 2007b, Limits to the sealing capacity of rock salt: A case study of the infra-Cambrian Ara Salt from the South Oman salt basin: Bulletin American Association Petroleum Geologists, v. 91, p. 1541-1557.

    Schreyer, W., and K. Abraham, 1976, Three-stage metamorphic history of a whiteschist from Sar e Sang, Afghanistan, as part of a former evaporite deposit: Contributions to Mineralogy & Petrology, v. 59, p. 111-130.

    Schreyer, W., K. Abraham, and H. Kulke, 1980, Natural sodium phlogopite coexisting with potassium phlogopite and sodian aluminian talc in a metamorphic evaporite sequence from Derrag, Tell Atlas, Algeria: Contributions to Mineralogy & Petrology, v. 74, p. 223-233.

    Scribano, V., S. Carbone, F. C. Manuella, M. Hovland, H. Rueslåtten, and H.-K. Johnsen, 2017, Origin of salt giants in abyssal serpentinite systems: International Journal of Earth Sciences, v. 106, p. 2595-2608.

    Shikazono, N., H. D. Holland, and R. F. Quirk, 1983, Anhydrite in kuroko deposits; mode of occurrence and depositional mechanisms: Economic Geology Monographs, v. 5, p. 329-344.

    Simoneit, B. R. T., 1994, Organic matter alteration and fluid migration in hydrothermal systems, in J. Parnell, ed., Geofluids: Origin, Migration and Evolution of fluids in Sedimentary Basins, Geological Society London, Special Publication No. 78, p. 261-274.

    Slack, J. F., M. R. Palmer, and B. P. J. Stevens, 1989, Boron isotope evidence for the involvement of non-marine evaporites in the origin of the Broken Hill ore deposits: Nature, v. 342, p. 913-916.

    Stewart, J. I., 1994, The role of evaporitic-shale sediment packages in the localisation of copper-gold deposits; Copper Canyon area, Cloncurry: Publication Series - Australasian Institute of Mining and Metallurgy, v. 5/94, p. 207-214.

    Svenningsen, O. M., 1994, Tectonic significance of meta-evaporitic magnesite and scapolite deposits in the Seve-Nappes, Sarek Mountains, Swedish Caledonides: Tectonophysics, v. 231, p. 33-44.

    Talbot, C. J., 2008, Hydrothermal salt--but how much?: Marine and Petroleum Geology, v. 25, p. 191-202.

    Tornos, F., J. M. Peter, R. Allen, and C. Conde, 2015, Controls on the siting and style of volcanogenic massive sulphide deposits: Ore Geology Reviews, v. 68, p. 142-163.

    Torres-Roldan, R. L., 1978, Scapolite-bearing and related calc-silicate layers from the Alpujarride Series, (Betic Cordilleras of southern Spain); a discussion on their origin and some comments: Geologische Rundschau, v. 67, p. 342-355.

    Trumbull, R. B., G. M. Garda, R. P. Xavier, J. A. D. Cavalcanti, and M. S. Codeço, 2018, Tourmaline in the Passagem de Mariana gold deposit (Brazil) revisited: major-element, trace-element and B-isotope constraints on metallogenesis: Mineralium Deposita.

    Van Driessche, A. E. S., J. M. García-Ruíz, K. Tsukamoto, L. D. Patiño-Lopez, and H. Satoh, 2011, Ultraslow growth rates of giant gypsum crystals: Proceedings of the National Academy of Sciences, v. 108, p. 15721.

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

    Webber, A. P., S. Roberts, B. J. Murton, and H. M. R. S., 2015, Geology, sulfide geochemistry and supercritical venting at the Beebe Hydrothermal Vent Field, Cayman Trough: Geochem. Geophys. Geosyst., v. 16, p. 2661–2678.

    Yardley, B. W. D., and J. T. Graham, 2002, The origins of salinity in metamorphic fluids: Geofluids, v. 2, p. 249-256.


    Non-solar thick salt masses: Part 1 - Mixing zone and cryogenic (freeze-dried) salts

    John Warren - Saturday, May 18, 2019


    Conceptually, what defines evaporite sediment is broad, within the general theme that the rock was precipitated initially during solar heating of a brine. Heating drives the loss of water as vapour and concentrates of the residual brine to elevated salinities where a suite of evaporite salts precipitate (Babel and Schreiber, 2014; Warren 2016). In a modern marine-fed brine pan the precipitative sequence evolves from carbonate through gypsum, to halite and on into the various bittern salts (Figure 1).

    But, evaporite salts are also highly reactive. Even in a simple single-pan evaporation scenario, crystallising salts tend to alter, backreact, replace, dissolve and reprecipitate. Then, as bedded sedimentary salts attain the subsurface and are exposed to increasing temperature and pressure, they continue to alter. Hydrated minerals, such as gypsum or carnallite transform into anhydrous phases such as anhydrite or sylvite. At the same time, thick halite beds tend to flow into halokinetic salt structures with a complete loss of depositional textures.

    Some authors restrict the term "evaporite" to sediments forged by evaporation and use "saline deposit" or "salt deposit" for units formed not only by evaporation but also by ongoing alteration, cooling, heating or salting out in hydrothermal and diagenetic settings. Hence, the term salt structure rather than evaporite structure to describe many subsurface features encompassing the outcomes of halokinesis (salt flow). Some authors have suggested other names for the various bedded salt rocks precipitated by mechanisms other than solar heating of a brine; however, these names (burial salt, hydrothermal salt, reactionite, mixing precipitate, thermalite, replacementite, etc.) are not in general use in the sedimentological community.

    In the subsurface sedimentary realm, most of the diagenetic processes described in preceding paragraphs have acted on a mass of salt first deposited by solar heating of brine and so are considered "true evaporites". This is so even though textures, total volumes of salts and mineralogies may have changed. Diagenetic processes acting on evaporite masses may also have transposed portions of the original salt mass into nearby saline cements within adjacent non-evaporite lithologies. These diagenetic or burial salt sediments lie outside the focus of this present series of articles, which will discuss mechanisms capable of precipitating large volumes of salts on or near the earth's surface without solar-driven heating of a brine (see Warren, 2016; Chapter 1 and 8 for a discussion of diagenetic or burial salts).

    Cryogenesis, brine-mixing and mantle-driven thermal processes are the main surface and shallow-subsurface processes capable of precipitating significant volumes of non-evaporite salts. Many of the salt bodies precipitated in this way have similar mineralogies to those found in evaporite successions. So this, and the next Salty Matters article focus on mechanisms and products created by non-evaporite precipitation. I will attempt to define criteria that allow their separation from "true" evaporites. In this first article, we focus on mechanisms of brine mixing and cryogenesis, in the second we shall look at salt masses crystallising from fluids created and driven by mantle heating and cooling.

    Some basic salt chemistry

    Before that, we need to discuss a few basic chemical properties that define a crystal's response to heating and cooling of an enclosing brine. The questions we must ask are; in the ambient conditions are we dealing with influencing a prograde or retrograde salt phase and does the particular salt of interest undergo congruent or incongruent dissolution?

    Prograde or retrograde salt?

    A prograde salt crystallises as a brine cools and dissolve as brine is heated. With prograde salts, the average kinetic energy of the molecules in brine increases with temperature. The increase in kinetic energy allows solvent molecules to more effectively break apart the solute molecules; hence solubility increases with temperature and decreases with cooling. Halite and sylvite are both prograde salts (Figure 2a). When a shallow surface brine rich in either NaCl or KCl at say 45-50°C cools overnight, it will precipitate halite or sylvite-carnallite on the pan floor. If it sinks into an underlying porous salt bed and cools, it will precipitate as intercrystalline cement. This is happening beneath the salt flats of Dabuxum Lake in the Qaidam Basin of China where intercrystalline carnallite fill pores in a halite bed. In the deeper crust, with ongoing magmatic heating, the solubility of sodium chloride increases to where just below its critical point (≈400°C), salt content in a NaCl brine could be as high as 40 weight percent (see next article).


    A retrograde (temperature-inverse) salt precipitates with increasing temperature and dissolves if it sits in a cooling saturated brine. Anhydrite is a retrograde salt and like all retrograde salts produces heat when dissolved in water (exothermic reaction). In an exothermic reaction the resulting additional heat shifts the equilibrium towards the reactants (Ca + SO4 --> CaSO4) (Figure 3b). In contrast, the hydrated form of CaSO4, gypsum, follows a prograde solubility curve, whereby in the temperature range where gypsum is the stable CaSO4 phase (<40-45°C) it is increasingly soluble with increasing temperature (Figure 3a). At higher temperatures (>45-50°C) both gypsum (where stable) and anhydrite possess retrograde solubility. Thus, when two anhydrite-saturated waters of different temperatures mix in the subsurface, the result is supersaturated brine with a propensity to precipitate anhydrite. A combination of retrograde solubility and brine mixing helps explain why anhydrite is a commonplace hydrothermal precipitate in mid-oceanic ridges and smoker chimneys (see next article).

    On land, beneath the immediate surface of Abu Dhabi and Saudi sabkhas, the heating of desert playa and sabkha waters, rising through the capillary zone, to temperatures above 35°C drives the precipitation of CaSO4 (as nodular anhydrite) in a fashion that Wood et al. (2002, 2005) terms a sabkha thermalite (see Warren, 2016 for sabkha discussion). A reverse thermal gradient (upward cooling) in the winter sabkha permits dissolution of some of the previously precipitated retrograde minerals. Because the negative thermal gradient is less in the winter than the positive thermal gradient of the summer, but with nearly the same water flux, the cooling water cannot dissolve all of the mineral mass that precipitated in the previous summer. Thus, there is a net accumulation of retrograde (thermalitic) anhydrite nodules and layers in the capillary zone of an Abu Dhabi sabkha (Wood et al. 2005 ).

    Brine mixing drives precipitation or dissolution of salts

    Van’t Hoff’s work in physical chemistry, which won him the first Nobel Prize in Chemistry, showed that salts would precipitate at chemically-suitable brine interfaces. At that interface, all that is needed is a mixing of waters of two saturation states, with respect to the mineral of interest (Figure 3a). When two waters that are saturated with a particular salt phase are mixed, the resulting solution can be undersaturated or supersaturated with respect to that particular phase (Figure 3a). The saturation state during mixing depends on the convex or concave shape of the solubility curve for the mineral phase of interest and the parameter of interest (ion concentration, temperature, salinity, etc.). The only requirement is that the solubility curve for that particular component is nonlinear.


    Raup (1970, 1982) in several experiments showed how halite and gypsum could be precipitated by the mixing of two seawater brines of differing salinity and densities. Figure 3b plots experimental results for the mixing of various low and high-density seawater brines and the resulting amount of gypsum precipitated (Raup, 1982). A similar plot can be drawn for the mixing of more saline NaCl (seawater) brines (Figure 4a, b; Raup, 1970).

    Hence blending brines with different temperatures or salinities can be an important salting-out mechanism in the calcium sulphate (gypsum/anhydrite) salt system where gypsum follows a nonlinear solubility trend, as do saturated brines in a halite-MgCl2 bittern system (Raup, 1970, 1982). Figure 2b shows the solubility curves of both gypsum and anhydrite plotted with respect to increasing temperature in pure water. For gypsum, it clearly shows that when two gypsum-saturated waters with different temperatures mix in its stability range then the resulting solution is undersaturated and so gypsum tends to dissolve (Upper left curve in Figure 2b). In contrast, gypsum drops out of solution when two brines of differing density mix, with the amount of gypsum dependent on the density contrast between the two brines (Figure 3b).

    Anhydritic solubility drives complex diagenetic effects when CaSO4-saturated brines mix via dispersion into adjacent less saline brines (Figure 2b). This happens in brine reflux systems where dense CaSO4-saturated brine plumes, derived at the surface at halite saturation, sink into and interact with less saline brines held in underlying or adjacent anhydritic carbonates. If the temperature remains near constant the tendency in this zone of dispersion or mixing is to dissolve gypsum, creating vugular porosity in the interval below or adjacent to a thick salt sequence. If temperature decreases and the brine plume cools, with little change in ionic proportions due to mixing, then the tendency is to precipitate anhydrite. It is not a simple system, with temperature and mixing processes pulling the brine chemistry in opposite directions.

    A similar salting-out of halite occurs when halite-saturated brine mix with either MgCl2 or CaCl2 brine (Figure 4a, b, respectively). In potash basins, MgCl2-saturated brines are created by the incongruent dissolution of carnallite, while CaCl2 brines can typify the brine products of basinal hydrothermal waters. In both cases, the result is a sparry, relatively inclusion-free halite cement. Depending or the location where this mixing occurred, the cement can be part of a bedded carnallite unit that is converting to sylvite, or it can be localised in a fracture fill or form an intergranular cement in a non-evaporite host lithology.

    Salt precipitation, driven by brine mixing, occurs in many stratified at-surface and shallow subsurface diagenetic interfaces in evaporitic settings. But, it is in the context of the mixing of deeply circulating mantle brines that it may be capable of precipitating significant volumes of salt and it in this context we shall discuss it further in the next article.

    Cryogenic salts

    Cryogenic brines and associated salts require temperatures at or below the freezing point of the liquid phase. These salts crystallise from a cold, near-freezing, residual brine as it concentrates via the loss of its liquid phase, which is converting/solidifying to ice (Figure 5a). As brine concentration increases, the freezing temperature decreases and minerals such as ikaite, hydrohalite, mirabilite, epsomite, potash bitterns and antarcticite can crystallise from the freezing brine (Figure 5a, b; Table 1; Warren, 2016; Chapter 12). Brine freezing ends when the phase cehemistry attains the eutectic point is reached. This is the point when all compounds (including H2O) pass to the solid state. Depending on the initial mineralization and compostion of the brine, the eutectic point is reached between -21 and -54 °C (Marion et al., 1999; Strakhov, 1970).

    Cryogenic concentration of seawater precipitates mirabilite at four times seawater salinity and hydrohalite at eight times. In contrast, evaporating seawater precipitates gypsum at 4-5 times the original concentration and halite and 10-11 times (Figure 6). Evaporative gypsum precipitation decreases the relative proportions of both Ca and SO4 in the brine, while cryogenic precipitation of mirabilite decreases the sulphate proportion and drives the inflexion of the Na cryogenic curve slightly earlier than Na inflexion created by the precipitation of evaporative halite. In both the freezing and the evaporation situations, the brine remains chloride dominant prior to bittern crystallisation. Freezing seawater becomes increasingly sulphate enriched to where sulphate levels exceed sodium around 20 times the initial concentration. Evaporating seawater remains a Na-Cl dominant brine until the bittern stage is reached around 60 times the initial concentration (Warren, 2016, Chapter 2).

    Mirabilite (NaSO4.10H2O) is one of several sodium sulphate salts and is stable in sulphate brines at temperatures lower than a few centigrade degrees (Figure 7a). Hence, it is a commonplace cold-climate lacustrine precipitate (Figure 8a-d; Table 1). Mirabilite beds are commercially exploited in colder climates; their latitudinal and altitudinal occurrences illustrate an interesting climatic dichotomy inherent to economic deposits of the various sodium sulphate salts. One sodium sulphate grouping of exploited deposits is characterised by mirabilite precipitated via brine freezing, as in the Great Salt Lake, Karabogazgol and Hedong (Yucheng) salt lake (Table 1 and illustrated in Figure 8a-d). The other sodium sulphate salt grouping is characterised by varying combinations of glauberite (CaSO4.Na2SO4)/bloedite-astrakanite (Na2SO4.MgSO4.4H2O) salts, which crystallised at higher temperatures via the evaporation of continental brines in saline groundwater sumps in warm to hot arid climates (as discussed in Warren, 2016, Chapter 12).

    The climatic dichotomy reflects the fact that sodium sulphate solubility in water changes as a nonlinear function of temperature (Figure 7a). Below 1.2°C, ice and mirabilite tend to precipitate as seawater or a sodium sulphate brine freezes. As the temperature increases above 0°C, increasing amounts of hydrous sodium sulphate (as the decahydrate, mirabilite) become soluble, while the anhydrous form (thenardite- NaSO4) becomes the precipitative phase in brines saturated with respect to the sodium sulphate. At 32.4°C in pure water, a transition point on the solubility curve is reached, whereby mirabilite melts in its water of crystallisation and thenardite crystallises. Presence of other dissolved salts changes the transition temperature and solubility characteristics of sodium sulphate due to the double salt effect.

    Cryogenesis in saline lacustrine sumps

    Every year, once the water temperatures drop below 5.5-6°C in late November in Karabogazgol, mirabilite precipitates as transparent crystals on the embayment bottom. The crystals are then transported by wind and wave action, especially during winter storms, into low dunes lining shore-zone (Figure 8b). By mid-March when the bay waters are heated to over 6°C, the mirabilite on the bay floor begins to redissolve. By July–August the entire precipitated mirabilite crop in the bay has redissolved. Historically, the period from November through March was a period of ‘‘harvesting’’ mirabilite on the bay shores. In summer with its arid climate, any remaining strandzone salt converts to thenardite (Na2SO4), and this too was gathered at the end of each summer.

    Below the floor of Karabogazgol are four NaSO4 beds that are likely cryogenic remnants from colder climatic periods over the last 10,000 years (Figure 9; Karpychev, 2007). Back then, large amounts of mirabilite formed each winter, much like today but, unlike today, the colder more humid glacial climate meant the bay was not as subject to summer desiccation and warming. Dense residual bottom brines were perennially ponded and so preserved a summer-halite sealing bed. This allowed the underlying mirabilite/epsomite winter precipitates to be preserved across the lake floor. During the following winter, the process was repeated as mirabilite/epsomite/halite couplets stacked one atop the other to create a future ore horizon. In time, the combination of groundwater and exposure, especially nearer the Gulf’s strandzone, converted most of the mirabilite, along with epsomite, to astrakanite, and then both phases to glauberite in the upper three beds. This explains the association of the richer glauberite zones with the lake edges (Figure 9a; Strakhov, 1970). Water of crystallisation released by the subsequent mirabilite to thenardite conversion, slightly diluted any strong residual brines, facilitating a dominant sodium-sulphate mineral and brine composition across the bay (Kurilenko et al., 1988).

    Winter mirabilite also crystallises cryogenically from Kuchuk Lake brines on the Kulunda Steppe, southwest of Novosibirsk, western Siberia (Kurilenko, 1997; Garrett, 2001. The lake area is 170 km2, and brine depth is around 3.2 metres. In 1938 that lake was estimated to contain some 540 million mt of equivalent sodium sulfate. Thick, glassy mirabilite occurs as two crystalline layers, the upper one is around 3 m thick, and both layers are pure containing <1% other soluble salts. In total, the crystalline mirabilite interval ranges up to 7 m thick, covers around 133 km2 and is overlain by a 0.05- to 2-m thick unconsolidated interval of mud and salt oozes. The ooze typically contains some 40.5% salts, 20.6% water, and 38.9% insolubles (including considerable gypsum). Much of the mirabilite in the upper ooze layer has been transformed into thenardite, with the previous water of crystallisation supply much of the dense brine held in the ooze, which also contains halite, glauberite, hydrohalite and epsomite.

    Mirabilite crystallises from the lake brine during the winter and cool summer evenings (volume estimated to be ≈ 340-580 thousand mt/yr of mirabilite). Then, during the warm summer months, some of it converts to thenardite. A limited amount of insoluble accumulates with the crystals, forming thin layers of mud with the thenardite. Brine in the lake has a 10-31% soluble salt content, depending upon the season and lake level. Every every three years, at the end of summer this brine is pumped to solar ponds to allow cryogenic mirabilite to crystallise during the autumn (a process similar to the production of mirabilite in the Canadian Salt Lake (Warren 2016, Chapter 12). Residual brine in the ponds is returned to the lake before winter sets in, and the ponds are harvested as needed for the production of sodium sulfate (Charykova et al., 1996). For most of a year, the lake's surface brine is a magnesium chloride water, but during the summer it changes to a sodium sulfate base, because of the dissolving of the underlying mirabilite, thenardite, and glauberite held in the lake floor oozes.

    Ebeity (Ebeyty) Lake, located 110 km west of Omsk, is another cryogenic salt lake with a cyclic pattern of mirabilite crystallising in the winter and having it dissolve in the summer (Garrett, 2001; Kolpakova et al., 2018). Brine concentration can reach 30-31% total salts by the end of summer and can begin to crystallise halite (which usually dissolves in the spring). Mirabilite deposition starts when the brine temperature in the lake is less than 18-19°C (which can be as early as August or September). At 0°C, 70% of the sodium sulfate has crystallised from the lake brine. At -10°C, 85% has been deposited, and at -15°C 98%. At -7°C some ice crystallizes with the mirabilite, and at -21.8°C hydrohalite forms. The lake does not freeze solid because of the insulating effect of surface layers of snow on top of floating mirabilite rafts, but brine temperatures of -23.5°C have been recorded (Strakhov, 1970). The winter deposit of mirabilite, with some hydrohalite, covers the entire lake bottom is 25-30 cm thick and is quite pure. Laboratory tests have shown the soluble salts in this mirabilite, including hydrohalite, can be almost completely removed (i.e., reduced to 0.08%) by a single stage of washing.

    Hydrohalite is a stable precipitate in a freezing brine only when the water temperature is below 0°C. In a NaCl–H2O system in the laboratory, hydrohalite is a stable phase that begins to precipitate cryogenically at temperatures below 0.12 °C, forming hydrohalite and a brine solution until it reaches the eutectic point for a solution saturated with NaCl at −21.1°C, where the remaining solution freezes (Figure 7b). Above 0.12 °C hydrohalite melts incongruently and decomposes to NaCl and a NaCl-saturated solution, losing 54.3% of its volume (Craig et al., 1974, 1975; Light et al., 2009). Hydrohalite (NaCl.2H2O) crystals have pseudo-hexagonal cross sections and are found in a number of modern cold saline lakes and springs (Table 1, Figure 8e-f).

    Cryogenesis in salty springs

    The mutual occurrence and downdip evolution of mirabilite/thenardite and hydrohalite/halite in brine spring encrustations (barrage structures) downdip of Stolz diapir on Axel Heiberg Island in the Canadian Archipelago illustrate the ephemeral nature of cryogenic salts on the Earths surface, even in extremely cold settings (Fox-Powell et al., 2018; Ward and Pollard, 2018). The halite-exposed core of the Stolz dome rises some 250 m above the adjacent flood plain, while the salt/hydrohalite deposit occurs autochthonously within a narrow, steep-sided tributary valley carved by a small stream fed mainly by perennial groundwater discharge emanating from the base of the diapir. The host valley begins abruptly at the spring outlet and is incised through surficial colluvial and glacial sediments into steeply dipping bedrock (shale). The Stolz diapir is the only diapir within the archipelago where the halite core is exposed, and the surface is extensively karstified with a suffusion cover, along with several large sinkholes and collapse structures.

    The cryogenically influenced part of the at-surface salt deposit is approximately 800 m long and is thickest at the spring outlet (≈4.0 m) and gradually thins down-valley until it fans out, creating a salt pan that extends 300 m into the Whitsunday River floodplain. The morphology of the deposit is characterised by a series of salty barrage structures that staircase down the valley until the dispersing and dissolving at the valley opening. The barrages are constructed of salts but morphologically resemble typical fluvial travertines and tufas and are predominantly curvilinear with a downstream convexity, particularly the larger dams in the upper valley.

    Sub-zero air temperatures persist at the site for at least ten months of the year, and the spring’s outlet temperature is relatively constant at −1.9°C ±0.1°C, confirming the presence of permafrost and so facilitating the precipitation of mirabilite and hydrohalite via freezing of spring waters (Figure 8f). In July air temperatures reach 5° - 6°C.

    Although initial precipitates in the permafrost zone of the spring outflow are cryogenic (mirabilite and hydrohalite) the presence of these cryogenic salts in the spring precipitates is ephemeral (Figure 10).

    The main body of a deposit is layered (Figure 4a, b) with alternating light and dark bands interpreted by Ward and Pollard (2018) as alternating periods of winter hydrohalite deposition and periods of summer pool drainage, when hydrohalite decomposes, and halite/thenardite sediment is deposited (Figure 11). As layers likely reflect an annual couplet cycle, then single winter accumulations can be as thin as a few millimetres in some parts of the deposit and as much as half a meter in others. It appears the accumulation phase ends as pools drain in early May corresponding with mean daily temperatures rising above 0°C. The darker layers are generated during summer as the hydrohalite decomposes leaving a granular halite crust with a veneer of fine clastic sediment transported into the deposit by wind, rain and runoff from adjacent slopes (during the spring snowmelt). The contact between the primary sediment and halite is abrupt and unconformable. The excavation of the deposit revealed numerous thin frozen layers (Figure 4c). These layers first appear ≈47cm below the surface, and overlying unfrozen material is considered analogous to an active layer in permafrost (Ward and Pollard, 2018). Samples of frozen salt collected from this layer and exposed to ambient air temperatures in summer reverted to a mixture of brine and halite grains (Figure 4d-e). It is not clear if these are residual hydrohalite layers preserved by thicker precipitate accumulations or if they represent secondary hydrohalite formation.

    The lack of mirabilite or thernadite within the upper portion of the stream/spring deposit is thought to be due to the low kinetic rates of precipitation for sulfate salts (Ward and Pollard, 2018). Below the halfway point, thernadite is present (Figure 8f). This is also reflected in the SO4 concentrations along the spring in winter, as the precipitation of mirabilite removes sulfate ions from solution beyond the halfway point. Based on the morphology of the crystals observed in the deposit, the low concentration of sulfate ions compared to chloride ions, as well as the extensive identification of halite in the XRD samples, the deposit is dominated by hydrohalite (not mirabilite), which decomposes to halite in the summer and at the same time the mirabilite that is present dehydrates to thenardite (Figure 11).

    Glacial and sea-ice cryogenesis

    Whenever polar seawater freezes, salts precipitate in the increasingly dense residual brines held in inclusions or fissures in the ice (Butler et al., 2016). Salts that are known to precipitate within the freezing brine include CaCO3.6H2O (ikaite), Na2SO4.10H2O (mirabilite), NaCl.2H2O (hydrohalite), KCl (sylvite), and MgCl2.12H2O (magnesium dichloride dodecahydrate), while hydrohalite is the most abundant salt to precipitate in sea ice (Light et al., 2009). In the seawater system, hydrohalite begins to precipitate at -22.9C, and further cooling results in additional precipitation until the source of Na is exhausted at the eutectic (Figure 7b).

    Salts do not only accumulate in sea ice. As dense brines and inclusion waters in flowing glacial ice sink into underlying rocks they can accumulate in ice sheet fissures at the base of the ice, or in load-induced fractures in the ice understory (Herut et al., 1990). Residual dense interstitial saline brines are found in pore waters extracted from deep cores sampling submarine sediments in McMurdo Sound, Antarctica (Frank et al., 2010). It seems that when ice sheets retreat, the at-surface cryogenic salts dissolve into a freshening at-surface hydrology, but dense hypersaline brines remain behind in deep fissures, held and preserved in the rock fractures (Starinsky and Katz, 2003).

    In the extreme setting of at-surface brine freezing in small saline depressions in the Dry Valleys of Antarctica, a solid form of calcium chloride, antarcticite, grows cryogenically in what is probably the most saline perennial natural water mass in the world (47% salinity in Don Juan Pond, Antarctica; Figure 8g-h; Horita et al., 2009).

    Mirabilite beds are known to be exposed atop ice floes of the Ross Ice Shelf near immediately down dip of the Hobbs glacier, on Cape Barne on Ross Island and in the vicinity of Cape Spirit, Black Island (Figure 12a; Brady and Batts, 1981). The bed is made up of relatively pure mirabilite that in places is more than a metre thick (Figure 12b). It is exposed in three coast-parallel ice pressure ridge systems and may not be continuous between the three sampling sites. According to Brady and Batts (1981), the mirabilite formed in response to a recent retreat of the Ross Ice Shelf that began some 840 years ago.

    The upper contact of the ice beneath the McMurdo mirabilite bed is not conformable as there are often small irregularities and undulations in its surface ranging from 2 to 6 cm high (Figure 12b). These undulations control the thickness of a silty sand interval (0-8cm thick) separating the ice from the mirabilite. The upper surface of the basal sediment layer defines a sharp conformable contact with the overlying mirabilite bed. This basal sediment layer is devoid of internal bedding and consists of 80% glacial flour mixed with sand and small pebbles. The sediment also contains many ice-crushed fragments of marine diatoms and sponge spicules (< 10 µm long).

    The overlying mirabilite bed is massive up to 1/2m thick and primarily made up of transparent cm-scale crystal clusters. Locally crystals can aggregate into large granules up to 95 mm across, that when exposed to air are coated by an anhydrous sodium sulphate powder rind Although there is no apparent bedding in the mirabilite bed, small pods and layered stringers of pebbly sand do occur. These are usually parallel or sub-parallel to the salt bed layer itself and vary in thickness from 0 to 12 cm. Broken shell fragments occur as rare isolated individual fragments in the salt. When the mirabilite is dissolved in distilled water, some fine mud and rare sand grains are recovered, as well as a perfectly preserved flora of non-marine diatoms.

    A lag of sediment, pebbles, cobbles, and boulders covers the mirabilite bed. The majority of the class are erratics, some of which are striated and come from the McMurdo alkaline volcanic province. There are also some erratics of gneiss, granite, and sandstone from continental suites. This lag is mostly overlain by a non-marine microbial mat (0-26 cm thick) but, in some places, the mat underlies or is mixed with the boulder lag. One mat sample collected 4 m above the level of the pool-and-channel systems on a pressure ridge at site 1, yielded a radiocarbon age of 870±70 years n.p. This single date cannot be used to date the whole mat since algae are still growing in pools in small depressions atop the salt deposit. The algal mat contains non-marine diatoms, but these are not as numerous as in nearby pools on the present-day ice shelf.

    Debenham (1920) suggested that the mirabilite he had observed on the Ross Ice Shelf was formed under the ice shelf by precipitation from brines. But it is unlikely that extensive linear pods and beds of friable salt could be brought directly to the surface by anchor ice; furthermore, the salt beds contain non-marine diatoms that indicate surface precipitation (Brady and Batts, 1981). Since non-marine diatoms have only settled in the salt itself, it would seem that the basal sediment layer was formed immediately after the injection of a sub-ice-shelf brine and before non-marine algal production in the brine pools.

    Hence, Brady and Batts (1981) conclude that mother brines first formed underneath an ice shelf from freezing sea-water. These brines were displaced to the ice-shelf surface by the weight of a large area of thick ice shelf as it grounded. Fine marine sediment carried in suspension by these brines settled to form an irregular thin discontinuous basal sediment layer containing marine diatoms. Mirabilite then crystallised cryogenically from this brine. During precipitation of the massive mirabilite beds, some non-marine diatoms, which can tolerate the high salt content of Antarctic saline lakes, were deposited with the salt. After the deposition of the mirabilite, massive non-marine algal production occurred, forming a thick irregular mat up to 26 cm thick on the mirabilite surface.

    Mirabilite beds lying on the Ross Island coast near Cape Barne and on the mainland near Hobbs Glacier likely formed in the same manner as those at Cape Spirit. That is, they were stranded on the coast as the ice shelf retreated to its present position in the south of McMurdo Sound during the last interglacial period.

    Extraterrestial cryogenic salts

    Cryogenic gypsum is released via ice ablation and is spread widely by katabatic winds across the circumpolar Martian dunefields (Table 2). Hydrated and calcium perchlorate cryogenic salts grow seasonally in soils of Mars and typify slope lineae on parts of the Marian surface Lineae activity is possibly tied to periodic release of liquid water in paces on the Martian surface. More than 3 billion years ago, evaporitic halite once precipitated in impact sumps on the surface of Mars.

    Cryogenesis explains the presence of hydrated magnesium sulphate salts in megapolygonal ice-crack fissures that crisscross the icy surfaces of Europa and Ganymede (moons of Jupiter) (Table 2; (Figure 13). The presence of hydrated magnesium sulphate and sodium carbonate salts indicates the presence of liquid salty oceans up to 100 km deep below an icy crust that is tens of kilometres thick (McCord et al., 1998; Craft et al., 2016. The fissures indicate an icy type of plate tectonics driven by strong tides in response to the varying gravitational pull of nearby Jupiter. Similar plumes of icy saline water escape from cracks and cryovolcanoes on the surface of Enceladus, an icy moon that circle Saturn. Spectral analysis shows the escaping plumes of Enceladus contain a variety of sodium and potassium salts (Postberg et al., 2011.

    Recognition of ancient terrestrial cryogenic salts

    Outside of relatively unaltered cool-temperature Quaternary lacustrine, permafrost and ice examples, as listed in Table 1, dehydration linked to the heating inherent to burial diagenesis will create mineralogical and textural difficulties in reliably interpreting the remains of ancient cryogenic salt beds. This is because all ionic salts forming in ambient surface conditions become metastable in the subsurface as they experience increased temperatures, pressures and evolving pore-fluid chemistries inherent to the diagenetic realm (Warren, 2016).

    Across all subsurface and re-exhumed ancient examples, the original low-temperature cryogenic salts (mirabilite, hydrohalite) will be long gone. Anhydrous salty remnants (thenardite, halite) may be preserved but are typically altered, replaced and dissolved. This makes it more challenging to assign a depositional setting to a cryogenic salt than it is to a typical evaporite.

    In their study of Miocene lacustrine thenardite in the Tajo Basin, Spain, Herrero et al. (2015) used three main criteria to suggest the conversion from a cold climate mirabilite precursor They were; 1) inclusion chemistry in the thenardite, 2) cooler climate mammal fauna synchronous with deposition of sodium sulphate salts and, 3) widespread dewatering structures tied to a burial transition from mirabilite to thenardite. A perusal of the depositional settings in the Quaternary examples listed in Table 1, underlines the conclusion that a glaciogenic association should be added as a fourth recognition criterion. Almost every case is either underlain or overlain by varying combinations of glacial till, glacial flour or glacial laminites with dropstones (Table 3).

    In terms of pre-Quaternary deposits of sodium sulphate salts, it was noted by Warren (2010, 2016), that the greater majority of economic deposits are Neogene sediments. Pre-Neogene cryogenic deposits are typically so diagenetically altered that no significant volumes of the original metastable NaSO4 cryogenic salts remain. Across all pre-Neogene greenhouse climate settings, that is across times that lack permanent polar ice sheets, there is little documentation of preserved volumes of cryogenic salts. Worldwide warmer temperatures mean it is unlikely there were extensive cryogenic salt beds. This leaves past periods in Earth history when the planet was in ice-house mode as possible times of cryogenic salt deposits were possible. As yet, no pre-Neogene cryogenic salt beds are known. Mirabilite/hydrohalite deposits have been inferred to be reasonable at tropical latitudes during the Neoproterozoic snowball period (Light et al. 2009; Carns et al., 2015)). Their presence has bee inferred to have increased albedo in tropical ice sublimation regions and so modify climate models, but as yet no evidence of their existence is known either then or in younger Ordovician or Permo-Carboniferous ice-age sediments. Likewise, the presence of glauberite in Permian lithologies is not diagnostic; glauberite forms from evaporating marine waters at times of MgSO4 -enriched oceans (Hardie 1985; Warren, 2016).

    Cryogenic salts are an extreme end-member of the concept of "the salt that was." Across past geological time, beyond Neogene remnants, even with isotopic and inclusion techniques, it is next to impossible to identify a cryogenic evaporite reliably.

    As W. Edwards Deming (Engineer and statistician, 1900-1993) once said "...Without data, you're just another person with an opinion."


    Alderman, S. S. J., 1985, Geology of the Owens Lake evaporite deposit, in B. C. Schreiber, and H. L. Harner, eds., Sixth international symposium on salt, v. 1, Salt Institute, VA, p. 75-83.

    Bąbel, M., and B. C. Schreiber, 2014, 9.17 - Geochemistry of Evaporites and Evolution of Seawater, in H. D. Holland, and K. K. Turekian, eds., Treatise on Geochemistry (Second Edition): Oxford, Elsevier, p. 483-560.

    Bowser, C. J., T. A. Rafter, and R. F. Black, 1970, Geochemical evidence for the origin of mirabilite deposits near Hobbs Glacier, Victoria Land, Antarctica: Mineralogical Society America Special Paper, v. 3, p. 261-272.

    Brady, H. T., and B. Batts, 1981, Large salt beds on the surface of the Ross ice shelf near Black Island, Antarctica: Journal of Glaciology, v. 27, p. 11-18.

    Braitsch, O., 1964, The temperature of evaporite formation, in A. E. M. Nairn, ed., Problems in palaeoclimatology: New York, Wiley, p. 479-490.

    Butler, B. M., S. Papadimitriou, A. Santoro, and H. Kennedy, 2016, Mirabilite solubility in equilibrium sea ice brines: Geochimica et Cosmochimica Acta, v. 182, p. 40-54.

    Charykova, M. V., 1996, The Hydrochemical Regime of Lake Kachuk: Water Resources, v. 23, p. 650-655.

    Craft, K. L., G. W. Patterson, R. P. Lowell, and L. Germanovich, 2016, Fracturing and flow: Investigations on the formation of shallow water sills on Europa: Icarus, v. 274, p. 297-313.

    Craig, J. R., J. F. Light, B. C. Parker, and M. G. Mudrey, 1975, Identification of hydrohalite: Antarctic Journal, v. 10, p. 178-179.

    Craig, J. R., B. L. Weand, and R. D. Fortner, 1974, Halite and Hydrohalite from Lake Bonney, Taylor Valley, Antarctica: Geology, v. 2, p. 389-390.

    Crétaux, J.-F., R. Létolle, and S. Calmant, 2009, Investigations on Aral Sea Regressions from Mirabilite Deposits and Remote Sensing: Aquatic Geochemistry, v. 15, p. 277-291.

    De Sanctis, M. C., A. Raponi, E. Ammannito, M. Ciarniello, M. J. Toplis, H. Y. McSween, J. C. Castillo-Rogez, B. L. Ehlmann, F. G. Carrozzo, S. Marchi, F. Tosi, F. Zambon, F. Capaccioni, M. T. Capria, S. Fonte, M. Formisano, A. Frigeri, M. Giardino, A. Longobardo, G. Magni, E. Palomba, L. A. McFadden, C. M. Pieters, R. Jaumann, P. Schenk, R. Mugnuolo, C. A. Raymond, and C. T. Russell, 2016, Bright carbonate deposits as evidence of aqueous alteration on (1) Ceres: Nature, v. 536, p. 54.

    Debenham, F., 1920, A new mode of transportation by ice: the raised marine muds of south Victoria Land (Antarctica): Quarterly Journal of the Geological Society of London, v. 75, p. 51-78.

    Fox-Powell, M. G., G. R. Osinski, M. Gunn, D. Applin, E. Cloutis, and C. R. Cousins, 2018, Low-Temperature Hydrated Salts on Axel Heiberg Island, Arctic Canada, as an Analogue for Europa. . Available from In 49th Lunar and Planetary Science Conference. Lunar and Planetary Institute, Houston. Abstract #2564.

    Frank, T. D., Z. Gui, and t. A. S. M. S. S. Team, 2010, Cryogenic origin for brine in the subsurface of southern McMurdo Sound, Antarctica: Geology, v. 38, p. 587-590.

    Garrett, D. E., 1998, Borates: Deposits, processing, properties and use: Amsterdam, Elsevier, 483 p.

    Garrett, D. E., 2001, Sodium sulfate: Handbook of deposits, processing, properties and uses: Amsterdam, Elsevier, 384 p.

    Grasby, S. E., I. Rod Smith, T. Bell, and D. L. Forbes, 2013, Cryogenic formation of brine and sedimentary mirabilite in submergent coastal lake basins, Canadian Arctic: Geochimica et Cosmochimica Acta, v. 110, p. 13-28.

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

    Herrero, M. J., J. I. Escavy, and B. C. Schreiber, 2015, Thenardite after mirabilite deposits as a cool climate indicator in the geological record: lower Miocene of central Spain: Clim. Past, v. 11, p. 1-13.

    Herut, B., A. Starinsky, A. Katz, and A. Bein, 1990, The role of seawater freezing in the formation of subsurface brines: Geochimica et Cosmochimica Acta, v. 54, p. 13-21.

    Horita, J., 2009, Isotopic Evolution of Saline Lakes in the Low-Latitude and Polar Regions: Aquatic Geochemistry, v. 15, p. 43-69.

    Hu, Y.-B., D. A. Wolf-Gladrow, G. S. Dieckmann, C. Völker, and G. Nehrke, 2014, A laboratory study of ikaite (CaCO36H2O) precipitation as a function of pH, salinity, temperature and phosphate concentration: Marine Chemistry, v. 162, p. 10-18.

    Hussmann, H., F. Sohl, and T. Spohn, 2006, Subsurface oceans and deep interiors of medium-sized outer planet satellites and large trans-neptunian objects: Icarus, v. 185, p. 258-273.

    Hynek, B. M., M. K. Osterloo, and K. S. Kierein-Young, 2015, Late-stage formation of Martian chloride salts through ponding and evaporation: Geology, v. 43, p. 787-790.

    Iess, L., D. J. Stevenson, M. Parisi, D. Hemingway, R. A. Jacobson, J. I. Lunine, F. Nimmo, J. W. Armstrong, S. W. Asmar, M. Ducci, and P. Tortora, 2014, The Gravity Field and Interior Structure of Enceladus: Science, v. 344, p. 78.

    James, N. P., G. M. Narbonne, R. W. Dalrymple, and T. K. Kyser, 2005, Glendonites in Neoproterozoic low-latitude, interglacial, sedimentary rocks, northwest Canada: Insights into the Cryogenian ocean and Precambrian cold-water carbonates Geology, v. 33, p. 9-12.

    Karpychev, Y., 2007, Variations in the sedimentation in Kara Bogaz Gol Bay related to sea level fluctuations during the Novocaspian time: Oceanology, v. 47, p. 857-864.

    Kolpakova, M., O. L. Gaskova, O. Naymushina, and S. K. Krivonogov, 2018, Ebeity lake, Russia: Chemical-organic and mineral composition of water and bottom sediments: v13BeCTv1H T0MCK0,O nonv1TeXHW-leCK0,O YHv1BeprnTeTa. v1H)Kv1Hv1pv1H, ,eopecypcoB., v. 329, p. 111-123.

    Kosarev, A., A. Kostianoy, and I. Zonn, 2009, Kara-Bogaz-Gol Bay: Physical and Chemical Evolution: Aquatic Geochemistry, v. 15, p. 223-236.

    Kurilenko, V. V., I. G. Ruday, and A. A. Shvarts, 1988, The origin and commercial exploitation of subsurface brines in the northern half of Kara-Bogaz-Gol: International Geology Reviews, v. 30, p. 1238-1245.

    Last, W. M., 2002, Geolimnology of salt lakes: Geosciences Journal, v. 6, p. 347-369.

    Lebedeva, M., O. Lopukhina, and N. Kalinina, 2008, Specificity of the chemical and mineralogical composition of salts in solonchak playas and lakes of the Kulunda steppe: Eurasian Soil Science, v. 41, p. 416-428.

    Light, B., R. E. Brandt, and S. G. Warren, 2009, Hydrohalite in cold sea ice: Laboratory observations of single crystals, surface accumulations, and migration rates under a temperature gradient, with application to “Snowball Earth”: Journal of Geophysical Research: Oceans, v. 114.

    Makhnach, A. A., 2008, Diagenetic gypsum, anhydrite and halite in non-evaporite deposits of Belarus: Baltica,, v. 21, p. 25-39.

    Marion, G. M., R. E. Farren, and A. J. Komrowski, 1999, Alternative pathways for seawater freezing: Cold Regions Science and Technology, v. 29, p. 259-266.

    McCaffrey, M. A., B. Lazar, and H. D. Holland, 1987, The evaporation path of seawater and the coprecipitation of Br(-) and K(+) with halite: Journal of Sedimentary Petrology, v. 57, p. 928-937.

    McCord, T. B., G. B. Hansen, F. P. Fanale, R. Carlson, W. , D. L. Matson, T. V. Johnson, W. D. Smythe, J. K. Crowley, P. D. Martin, A. Ocampo, C. A. Hibbitts, and J. C. Granahan, 1998, Salts on Europa's surface detected by Galileo's near infrared mapping Spectrometer: Science, v. 280, p. 1242-1245.

    Muessig, S., 1958, First known occurrence of inyoite in a playa at Laguna Salinas: Am. Mineral., v. 43, p. 1144–1147.

    Nelson, K. H., and T. G. Thompson, 1954, Deposition of salts from seawater by frigid concentration: Journal Marine Reseearch, v. 13, p. 166-182.

    Nikolaevsky, A. P., 1938, The winter minerals of the Baskunchak salt lake: Priroda, v. 1, p. 86-93.

    Ojha, L., M. B. Wilhelm, S. L. Murchie, A. S. McEwen, J. J. Wray, J. Hanley, M. Masse, and M. Chojnacki, 2015, Spectral evidence for hydrated salts in recurring slope lineae on Mars: Nature Geosci, v. 8, p. 829-832.

    Peterson, R. C., and R. Y. Wang, 2006, Crystal molds on Mars: Melting of a possible new mineral species to create Martian chaotic terrain: Geology, v. 34, p. 957-960.

    Postberg, F., J. Schmidt, J. Hillier, S. Kempf, and R. Srama, 2011, A salt-water reservoir as the source of a compositionally stratified plume on Enceladus: Nature, v. 474, p. 620.

    Raup, O. B., 1970, Brine mixing - an additional mechanism for formation of basin evaporites: Bulletin American Association of Petroleum Geologists, v. 54, p. 2246-2259.

    Raup, O. B., 1982, Gypsum precipitation by mixing seawater brines: Bulletin American Association of Petroleum Geologists, v. 66, p. 363-367.

    Roberts, S. M., R. J. Spencer, W. B. Yang, and H. R. Krouse, 1997, Deciphering some unique paleotemperature indicators in halite-bearing saline lake deposits from Death Valley, California, USA: Journal of Paleolimnology, v. 17, p. 101-130.

    Ruesch, O., T. Platz, P. Schenk, L. A. McFadden, J. C. Castillo-Rogez, L. C. Quick, S. Byrne, F. Preusker, D. P. O’Brien, N. Schmedemann, D. A. Williams, J. Y. Li, M. T. Bland, H. Hiesinger, T. Kneissl, A. Neesemann, M. Schaefer, J. H. Pasckert, B. E. Schmidt, D. L. Buczkowski, M. V. Sykes, A. Nathues, T. Roatsch, M. Hoffmann, C. A. Raymond, and C. T. Russell, 2016, Cryovolcanism on Ceres: Science, v. 353, p. aaf4286.

    Smith, G. I., 1979, Subsurface stratigraphy and geochemistry of Late Quaternary evaporites, Searles Lake, California: US Geological Survey, Professional Paper, v. 1043, p. 130 pp.

    Starinsky, A., and A. Katz, 2003, The formation of natural cryogenic brines: Geochimica et Cosmochimica Acta, v. 67, p. 1475-1484.

    Strakhov, N. M., 1970, Principles of Lithogenesis (Reviews of USSR Sodium Sulfate Deposits): New York, Plenum Publishing.

    Vance, S., M. Bouffard, M. Choukroun, and C. Sotin, 2014, Ganymede's internal structure including thermodynamics of magnesium sulfate oceans in contact with ice: Planetary and Space Science, v. 96, p. 62-70.

    Ward, M. K., and W. H. Pollard, 2018, A hydrohalite spring deposit in the Canadian high Arctic: A potential Mars analogue: Earth and Planetary Science Letters, v. 504, p. 126-138.

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

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

    Wood, W. W., W. E. Sanford, and A. R. S. Al Habshi, 2002, Source of solutes to the coastal sabkha of Abu Dhabi: Geological Society of America Bulletin, v. 114, p. 259-268.

    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.

    Xiyu, Z., 1984, Distribution characteristics of boron and lithium in brine of Zhacang Caka salt lake, Xizang (Tibet), China: Chinese Journal of Oceanology and Limnology, v. 2, p. 218-227.

    Zavialov, P., A. Ni, T. Kudyshkin, D. Ishniyazov, I. Tomashevskaya, and D. Mukhamedzhanova, 2009, Ongoing Changes of Ionic Composition and Dissolved Gases in the Aral Sea: Aquatic Geochemistry, v. 15, p. 263-275.

    Zheng, M., 1997, An introduction to saline lakes on the Qinghai-Tibet, Plateau: Monographiae Biologicae, v. 76, Springer, 328 p.

    Zheng, M., T. Jiayou, L. Junying, and Z. Fasheng, 1993, Chinese saline lakes: Hydrobiologia, v. 267, p. 23-36.

    Zheng, M., and X. Liu, 2009, Hydrochemistry of Salt Lakes of the Qinghai-Tibet Plateau, China: Aquatic Geochemistry, v. 15, p. 293-320.


    Recent Posts


    NPHI log Neoproterozoic Oxygenation Event York (Whitehall) Mine base metal snake-skin chert Catalayud wireline log interpretation Gamma log silicified anhydrite nodules McMurdo Sound hydrothermal potash SO2 Beebe hydrothermal field collapse doline Clayton Valley playa: halokinetic Weeks Island salt mine sedimentary copper carnallitite salt mine supercritical phase well blowout Dallol saltpan geohazard lithium battery Mesoproterozoic Kalush Potash Sulphate of potash Belle Isle salt mine phreatic evaporite bischofite H2S jadarite Koppen climate Neutron Log Europe SOP gassy salt Deep methanotrophic symbionts lot's wife Magdalen's Road circum-Atlantic Salt Basins HYC Pb-Zn cryogenic spring salts salt leakage, dihedral angle, halite, halokinesis, salt flow, stable isotope CO2 Dead Sea saltworks gem organic matter Patience Lake member deep seafloor hypersaline anoxic basin silica solubility tachyhydrite doline CO2: albedo Phaneozoic climate salt seal Hell Kettle 13C enrichment cauliflower chert Precambrian evaporites sulphate vestimentiferan siboglinids sulphur stevensite Prairie Evaporite eolian transport nuclear waste storage CaCl2 brine antarcticite MOP deep meteoric potash High Magadi beds seawater evolution natural geohazard vadose zone African rift valley lakes 18O enrichment Ganymede Koeppen Climate crocodile skin chert water in modern-day Mars supercontinent Bathymodiolus childressi authigenic silica MVT deposit Schoenite palygorskite Salar de Atacama brine lake edge K2O from Gamma Log lithium brine zeolite Realmonte potash mass die-back salt trade halophile allo-suture Zaragoza well logs in evaporites chert Mixing zone NaSO4 salts Warrawoona Group Mega-monsoon astrakanite extrasalt salt periphery Mulhouse Basin marine brine methanogenesis Mars sodium silicate MgSO4 depleted rockburst Quaternary climate Prograde salt End-Triassic ancient climate hydrogen Deep seafloor hypersaline anoxic lake intrasalt hydrological indicator Ripon Turkmenistan GR log alkaline lake anthropogenically enhanced salt dissolution Calyptogena ponderosa water on Mars Stebnik Potash Large Igneous Magmatic Province anthropogenic potash salt ablation breccia sepiolite Kara bogaz gol evaporite-metal association Hadley cell: brine evolution Red Sea evaporite End-Cretaceous black salt capillary zone LIP Messinian anomalous salt zones venice cryogenic salt white smokers halotolerant waste storage in salt cavity carbon oxygen isotope cross plots Density log source rock basinwide evaporite Neoproterozoic trona lazurite Jefferson Island salt mine non solar heating salts freefight lake Badenian Danakhil Depression, Afar Lake Magadi retrograde salt saline giant recurring slope lines (RSL) dihedral angle hydrothermal karst phreatomagmatic explosion MgSO4 enriched lapis lazuli Zabuye Lake Muriate of potash extraterrestrial salt salt suture carbon cycle namakier Dead Sea caves nitrogen supercritical halite Musley potash North Pole Boulby Mine hydrothermal halite oil gusher causes of major extinction events Stolz diapir hectorite phreatic explosion salting-out evaporite-hydrocarbon association gas outburst meta-evaporite brine pan Pangaea Ceres mine stability Ure Terrace Karabogazgol ozone depletion Lamellibrachia luymesi hydrohalite Enceladus DHAL salt karst seal capacity Dead Sea karst collapse potash ore price evaporite karst vanished evaporite endosymbiosis Archean Lake Peigneur salt tectonics Proterozoic lunette saline clay epsomite well log interpretation Platform evaporite 13C Noril'sk Nickel halogenated hydrocarbon Corocoro copper Lop Nor sinkhole climate control on salt Atlantis II Deep DHAB evaporite dissolution Thiotrphic symbionts perchlorate Hadley Cell bedded potash mummifiction Lomagundi Event subsidence basin flowing salt magadiite Crescent potash dissolution collapse doline methane gypsum dune Hyperarid intersalt Evaporite-source rock association potash Pilbara End-Permian RHOB 18O solar concentrator pans Great Salt Lake halite-hosted cave Ethiopia Sumo potash ore Belle Plain Member Stebnyk potash kainitite causes of glaciation halocarbon hydrothermal anhydrite SedEx lithium carbonate knistersalz gas in salt Five Island salt dome trend sinjarite Ingebright Lake Seepiophila jonesi Paleoproterozoic Oxygenation Event Lop Nur sulfur auto-suture dark salt solikamsk 2 mirabilite halite nacholite McArthur River Pb-Zn sulfate blowout